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Exhumation and cooling history of the Middle Eocene Anaconda metamorphic core complex, western Montana


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mTtmg 36 24 24 89 35 80 68 49 80 75 35 22 31 19 72 20 30 20 34 30 47 40 32 37 17 19 50 20 19 5 35 32 89 82 56 21 5 82 84 85 60 77 50 37 76 36 60 76 84 75 53 55 48 47 69 80 63 60 67 55 85 81 85 85 50 65 54 55 29 20 50 38 24 28 65 54 20 68 49 24 1 1 1 1 15 14 75 20 28 27 70 19 34 20 15 14 21 20 45 47 55 23 24 21 28 40 25 4 14 25 15 44 6 12 8 0 10 18 34 69 25 25 59 10 45 50 25 24 60 42 25 65 70 52 34 14 44 55 32 48 45 20 52 49 20 16 80 40 34 30 42 89 86 45 79 79 87 84 45 50 27 28 57 43 36 74 5 50 10 33 5 67 67 10 35 30 18 20 23 17 40 25 17 27 17 5 32 38 5 1 20 10 22 16 14 28 25 27 16 46 33 3 15 18 19 20 15 23 42 46 22 33 13 13 30 15 23 33 28 38 38 60 78 45 45 42 55 45 32 70 5 0 80 38 71 53 52 24 50 49 48 40 16 35 40 13 56 54 24 46 74 24 20 30 65 47 24 15 20 44 44 90 25 25 15 29 20 22 52 43 66 38 69 40 36 81 51 51 77 37 Qu Qu Qu Qu Qu Qu Qu Qu Qu mTbg mTtmg Ttmg Kqd Kqd Kqd Yr Yr Kqd Yr Yr YLb? Yr Yh Yh Yh Yh Yh Yh Yh Yh Ygr Ttmg Ttmg Ygr Ygr Ygr Ygr Ttmg Kgd Kgd Tgd Tgd Tgd Tbg Tpg mq mq Tbg Tgd Tgd Ymg Ymg Ymg Ymg Ymg Ymg Csh Ch Cf Ch Csh Cf Csh Csh Cf Cf Csh Cf Csh Ch Kqd Cf Csh Ch tightly folded, attenuated Cambrian section Qu Qu Qu Qu Qu Qu Cf Cf Csh mTbg mTgd Tgd Ch Ygr Cf Ch Csh Ttmg Cf Ch Ch Ch Ch Ymg? Ch Kqd Ch Ch Ygr diabase sills T ac 70 Kgd YLb? Ygr GEOLOGIC MAP OF THE FOOTW ALL OF THE ANACONDA MET AMORPHIC CORE COMPLEX, NOR THERN ANACONDA-PINTLAR RANGE, WESTERN MONT ANA BY W ARREN C. GRICE 2006 Geology by W arren C. Grice (surveyed in summer 2004-2005) with the exception of the Mt. Haggin area geology which is taken from mapping of Heise (1983) and Kalakay et al. (unpublished mapping). Limited structural data and some unit contacts were taken from mapping of Emmons and Calkins (1913) and Lonn et al. (2004). Base topography from USGS 1:24000 scale Storm Lake, Mount Evans,and Mount Haggin 7.5 minute qradrangles. References cited: (1) Emmons, W H. and Calkins, F C., 1913, Geology and ore deposits of the Phillipsburg quadrangle, Montana: U.S. Geological Survey Professional Paper 78, 271 p. (2) Heise, B. A., 1983, Structural geology of the Mt. Haggin area, Deer Lodge country Montana. Unpublished Masters Thesis, University of Montana, Missoula, Montana, 77 p. (3) Lonn, J. D., McDonald, C., Lewis, R. S., Kalakay T J., O'Neill, J. M., Berg, R. B., and Hargrave, P ., 2003, Preliminary geologic map of the Philipsburg 30'x 60' quadrangle, western Montana: Montana Bureau of Mines and Geology Open file Report MBMG 483, scale: 1:100,000. Greenschist facies mylonite shear zone Upper-amphibolite facies ductile shear zone (The Lake of the Isle shear zone) Migmatitic pelitic schist and gneiss Normal fault Low-angle normal detachment fault Antiform Explanation of Map Symbols Thrust fault 32 38 21 89 Metamorphic foliation w/ dip Magmatic foliation w/ dip Metamorphic foliation w/ dip and trend of lineation V ertical metamorphic foliation 14 25 70 Joint w/ dip Mesoscopic fold w/ trend of hinge Cleavage w/ dip MAP SCALE 1:24000 1 2 km North 1 0 Contour interval is 40 feet Map projection: 1927 NAD (North American Datum) Declinatin is 16 East Sample locality Description of Mapped Units Yg Yr Yh Ymg Cf Csh Ch Kgd Kqd Ttmg Tbg Tgd Tpg T ac Qu Quaternary undivided glacial moraines and alluvium Anaconda Beds sediments and volcaniclastics Dacitic dike T wo-mica porphoritic granite Biotite granodiorite (some mylonitic) Biotite granite (often mylonitic, some porphoritic) T wo-mica granite (often mylonitic) Hornblende, biotite granodiorite (Storm Lake stock) Quartz diorite (sill in central and east is deformed) Hasmark formation white, coarse marble Silver Hill formation mica schist Flathead quartzite white to pink, quartzite Missoula Group green calc-silicate and biotite quartzite gneisses Helena formation calc-silicate gneiss Ravalli Group biotite quartzite gneiss and schist Greyson formation grt+bt+sill schist and gneiss (locally migmatitic) mq Mylonitic quartzite white, mylonitic quartzite (Flathead or Niehart equilavent?) Td Uncorrelated Units Ylb Biotite quartzite gneiss and schist (Ravalli or Greyson equilavent?) 113 15' 46 7.5' Region not mapped 113 Region not mapped



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EXHUMATION AND COOLING HIST ORY OF THE MIDDLE EOCENE ANACONDA METAMORPHIC CORE COMPLEX, WESTERN MONTANA By WARREN CALHOUN GRICE, JR. A THESIS PRESENTED TO THE GRADUATE SCHOOL OF THE UNIVERSITY OF FLOR IDA IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE UNIVERSITY OF FLORIDA 2006

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Copyright 2006 by WARREN CALHOUN GRICE, JR.

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Dedication: To my family for thei r never ending support and inspiration.

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iv ACKNOWLEDGMENTS I would like to thank David Foster, my advisor, for hi s guidance and patience while working with me on this study over the past few years. I thank Tom Kalakay for introducing me to geology and for instilling in me great passion and motivation for the subject; I also thank Kalakay for his essent ial guidance in the fi eld during the 2004 field season. I am greatly appreciative of Jim V ogl, Phil Neuoff, Paul Mueller, Matt Smith, and Mike Perfit for their assi stance and guidance with seve ral aspects of the study. I thank George Kamenov, Sam Coyner, and Tom B easley who helped make the analytical aspects of the study possible. I thank Dan Go rman for help with sample preparation. I would also like to thank Michaela Speirs Heather Bleick, Sykl ar Pauli, and Mike McTeauge for their assistance in the field; I also thank McTeague for sharing my joy for geology over the past several years and for many interest ing geologic discussions. Financial support for the field aspect of this study was supplied by grants obtained by author from the Tobacco Root Geological So ciety, the Geological Society of America, and the Belt Association. Funding for the analytical portion of the study was supplied by a USGS EDMAP grant obtained by Tom Ka lakay and by grants obtained by David Foster. The study would not have been possible wit hout this funding.

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v TABLE OF CONTENTS page ACKNOWLEDGMENTS.................................................................................................iv LIST OF TABLES.............................................................................................................ix LIST OF FIGURES.............................................................................................................x ABSTRACT.....................................................................................................................xi v CHAPTERS 1 INTRODUCTION........................................................................................................1 2 REGIONAL GEOL OGIC BACKGROUND...............................................................9 Pre-Mesozoic History...................................................................................................9 Mesozoic History........................................................................................................13 Crustal Shortening along the Lewis and Clark Line during the Late Cretaceous........................................................................................................14 Late Cretaceous Magmatism...............................................................................16 High Grade Metamorphism during the Cretaceous.............................................19 Eocene History............................................................................................................20 Regional Transtension and Large-scal e Crustal Extension along the Lewis and Clark Line..................................................................................................21 3 THE ANACONDA METAMO RPHIC CORE COMPLEX......................................23 Structural-metamorphic Domains of the Anaconda Metamorphic Core Complex:...23 The Lower Plate..................................................................................................23 The Detachment Fault Zone................................................................................29 The Upper Plate...................................................................................................33 Description of the Lake of the Isle Shear Zone..........................................................35 Description and Distribution of Meta morphic and Structural fabrics.................36 Metamorphic foliations and gneissic banding..............................................36 Mesoscopic-scale folds................................................................................41 Mesoscopic-scale boudins............................................................................42 Shear sense...................................................................................................45 Description and Distribution of Me tamorphic Phase Assemblages and Textures............................................................................................................45

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vi Metamorphic phase assemblages an d textures in the meta-Greyson Formation..................................................................................................46 Relict and fresh kyanite in the upper Meta-Ravalli Group..........................50 Spatial Relationship Between the Lake of the Isle Shear Zone and Late Cretaceous Intrusions..............................................................................................53 4 U-PB ZIRCON GE OCHRONOLOGY......................................................................58 Purpose and Strategy..................................................................................................58 Relevant Previous UPb Zircon Geochronology........................................................59 U-Pb Zircon Geochronology Results..........................................................................59 WG04-114 (Storm Lake Stock Granodiorite).....................................................60 Ug-1 (The Deformed Quartz Diorite Sill)...........................................................61 WG05-02 (Leucosome from the Meta-Greyson Paragneiss)..............................70 5 THERMOBAROMETRY..........................................................................................76 Purpose and Strategy..................................................................................................76 ME-231 (Migmatitic Meta-Greyson Formation Paragneiss)......................................77 Relevant Previous Pressure -temperature Constraints.................................................78 Thermobarometry Results..........................................................................................80 Electron Microprobe Analyses............................................................................80 Garnet...........................................................................................................80 Biotite...........................................................................................................86 Plagioclase....................................................................................................86 Cordierite......................................................................................................86 Mineral End Member Activity Calculations made using AX.............................87 Pressure-temperature Estimates using THERMOCALC....................................87 Pressure-temperature Estimates using Geothermobarometry (GTB)..................93 6 40AR/39AR THERMOCHRONOLOGY.....................................................................95 Purpose and Strategy..................................................................................................95 Previous Thermochronology......................................................................................96 40Ar/39Ar Thermochronology Results ......................................................................104 Argon Closure Temperature..............................................................................105 Biotite................................................................................................................122 Muscovite..........................................................................................................125 Hornblende........................................................................................................128 K-feldspar..........................................................................................................131 40Ar/39Ar Thermochronology From Outside the Lower Plate Transect...................133 7 DISCUSSION...........................................................................................................136 Origin of the Lake of the Isle Shear Zone................................................................136 Age Constraints.................................................................................................137 Pressure-temperature History............................................................................140 Kinematic Interpretation....................................................................................144

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vii Strain localization and ductile attenu ation of metasedimentary strata in the LISZ..................................................................................................145 Strain classification....................................................................................146 Structural Interpretation.............................................................................146 Eocene Exhumation and Cooling History of the Anaconda Metamorphic Core Complex defined by 40Ar/39Ar Thermochronology..............................................147 Lower Plate Cooling History.............................................................................148 Constraints on the Timing of the Onset of Extension.......................................154 Constraints on the Duration of Extension.........................................................159 Constraints on the Deta chment Slip Rate..........................................................161 Constraints on the Original Detachment Geometry..........................................169 Magnitude of Offset on the Detachment...........................................................177 8 CONCLUSIONS......................................................................................................183 Age and Pressure-temperature Constrai nts on High Grade Metamorphism in the ACC lower plate: Origin of th e Lake of the Isle Shear Zone...............................183 Exhumation and Cooling History of th e Middle Eocene Anaconda Metamorphic Core Complex:......................................................................................................183 Regional Tectonic Context.......................................................................................188 APPENDIX A METHODOLOGIES................................................................................................191 Field Mapping and Sampling Methods.....................................................................191 40Ar/39Ar Thermochronology Methods.....................................................................192 Thermobarometry.....................................................................................................195 U-Pb Geochronology................................................................................................196 Major and Trace Elem ent Geochemistry..................................................................203 B ACC LOWER PLATE LITHOLOGI C UNIT DESCRIPTIONS AND DISTRIBUTION .....................................................................................................205 General Statement.....................................................................................................205 Belt Supergroup-equivalent Metasedimentary Strata...............................................206 Middle Cambrian-equivalent Metasedimentary Strata.............................................210 Late Cretaceous Intrusions.......................................................................................212 The Storm Lake Stock.......................................................................................212 The Deformed Quartz Diorite Sill.....................................................................216 Eocene Intrusions......................................................................................................217 C DESCRIPTION OF SAMPLE ME-231...................................................................224 D MAJOR AND TRACE ELEMENT GEOCHEMISTRY.........................................229 Purpose.....................................................................................................................229 Major and Trace Element Analytical Results...........................................................230

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viii Major Element Analyses...................................................................................231 Trace Element Analyses....................................................................................231 E THERMOCALC AND AX OUTPUT FILES..........................................................240 F GEOLOGIC MAP OF THE ANACONDA METAMORPHIC CORE COMPLEX LOWER PLATE.......................................................................................................245 REFERENCES CITED....................................................................................................247 BIOGRAPHICAL SKETCH...........................................................................................261

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ix LIST OF TABLES Table page 3-1. Description and distribution of ke y mineral phases and textures in the metamorphosed Greyson Formation........................................................................54 4-1. U-Pb LA-MC-ICP-MS anal ytical results for WG04-114........................................62 4-2. U-Pb LA-MC-ICP-MS analytical results for Ug-1..................................................66 4-3. U-Pb LA-MC-ICP-MS anal ytical results for WG05-02..........................................68 5-1. Results from electron mi croprobe mineral analyses................................................81 5-2. Averaged ME-231 electron microprobe analyses used in thermobarometry...........88 6-1. Summary of relevant pr evious thermochronology...................................................98 6-2. Summary of 40Ar/39Ar thermochronology from the Anaconda metamorphic core complex..................................................................................................................106 D-1. Results from major and trace element analyses.....................................................232

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x LIST OF FIGURES Figure page 1-1. Regional geologic map of the North American Cordillera. Emphasis is placed on major igneous and tectonic features......................................................................2 1-2. Tectonic map of northwestern United States and southern Canada showing major structures related to Tertiary extension........................................................... 4 2-1. Regional geologic map of the major Archean basement provinces and Proterozoic suture zones of the North Am erican Cordillera in the western United States and southern Canada......................................................................................10 2-2. Simplified tectonic map of Belt Basin in the Mesoproterozoic. Map shows the major normal faults active during depos ition of the Belt-Pu rcell Supergroup........12 2-3. Tectonic map of the North American Cordillera of western Montana, northern Idaho, northeastern Washington, and s outhern Canada during the Late Cretaceous................................................................................................................15 2-4. Geologic map of western Montana, south of the Lewis and Clark line...................18 3-1. Geological Map of the Anaconda metamo rphic core complex (ACC) in western Montana....................................................................................................................25 3-2. Geologic cross section of the Anac onda metamorphic core complex, western Montana....................................................................................................................26 3-3. Photomosaic of the Anaconda metamorphic core complex.....................................32 3-4. Greenschist facies mylonites....................................................................................33 3-5. Outcrop photos of late li stric-shaped brittle normal faults in the greenschist facies mylonite zone.................................................................................................34 3-6. Outcrop photograph showing the str ong gneissic banding of the metamorphosed Greyson Formation in the LISZ...............................................................................39 3-7. Outcrop photographs of gneissic banding commonly observed in the LISZ...........40 3-8. Photomicrograph of the Missoula Groupequivalent calc-sil icate paragneiss.........41

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xi 3-9. Outcrop photograph of the deformed quartz diorite sill in the center of the LISZ.........................................................................................................................43 3-10. Outcrop photographs of mesoscopicscale folds found in Belt equivalent metasedimentary strata deformed in the LISZ.........................................................44 3-11. Outcrop photographs of mesoscopicscale folds found in Belt-equivalent metasedimentary strata deformed in the LISZ.........................................................46 3-12. Outcrop photographs of me soscopic boudins in the LISZ.......................................47 3-13. Outcrop photographs of shear sense in dicators in the meta-Greyson Formation deformed in the LISZ...............................................................................................49 3-14. Outcrop photographs of meta-Greyson migmatitic paragneiss showing granitic leucosome commonly found in pressure shadows between quartzite boudins........51 3-15. Photomicrographs of thin sections of the meta-Greyson mi gmatitic paragneiss exhibiting granitic leuc osome (sample ME-231).....................................................52 3-16. Simplified geologic sketch map of the ACC lower plate exposed in the current study area showing the metamorphic mi neral/textural zones 1, 2, and 3 mapped in the meta-Greyson Formation adjacent to the deformed quartz diorite sill...........55 3-17. Kyanite pseudomorphs in the uppe r meta-Ravalli Group. A) Outcrop photograph of kyanite pseudomorphs......................................................................56 3-18. Photomicrograph showing fresh kyanite in the upper meta-Ravalli Group.............57 4-1. Tera-Wasserburg plot and 206Pb / 238U weighted mean zircon age plot for Storm Lake Stock granodior ite sample WG04-114............................................................70 4-2. Tera-Wasserburg plot and 206Pb / 238U weighted mean zircon age plot for the deformed quartz diorite sill sample Ug-1.................................................................71 4-3. Conventional U-Pb concordia plots fo r zircons for leucosome sample WG05-02 collected from the meta-Greyson migma titic paragneiss in the central LISZ..........74 4-4. 207Pb/206Pb age density plot for the WG05-02 leucosome zircons. 207Pb/206Pb ages range from ~1525-3004 Ma.............................................................................75 5-1. Phase diagrams showing PT estim ates for ME-231, migmatitic Greyson paragneiss from the Lake of the Isle shear zone......................................................91 5-2. Phase diagram showing phase equilibr ia used by THERMOCALC to calculate the average PT estimate for sample ME-231...........................................................93 6-1. Geologic map of Anaconda metamo rphic core complex (ACC), western Montana and vicinity showing loca tion of previous thermochronology................102

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xii 6-2. Geologic sketch map showing mineral cooling ages from samples collected along the ACC lower plate transect using 40Ar39Ar thermochronology.................109 6-3. 40Ar/39Ar mineral age spectra obtained for samples collected from the Anaconda metamorphic core complex using 40Ar/39Ar thermochronology............................110 7-1. Phase diagram showing the Late Cretaceous PT history of the Lake of the Isle shear zone (LISZ)...................................................................................................142 7-2. Geologic sketch map showing the mi neral cooling ages obtained by 40Ar39Ar thermochronology from samples collected along the ACC lower plate transect in this study................................................................................................................149 7-3. Mica cooling age contour map construc ted from biotite and muscovite cooling ages obtained 40Ar39Ar thermochronology from samples collected along the ACC lower plate transe ct in this study...................................................................150 7-4. Temperature-time cooling diagram s howing the cooling history of the ACC lower plate..............................................................................................................151 7-5. Age vs. distance in slip direction diag ram and lower plate transect sketch map...159 7-6. Slip rate calculations..............................................................................................166 7-7. Paleoisotherm contour map. Paleoisother ms refer to estimated temperatures of the ACC lower plate directly beneath the de tachment at the onset of extension at ~53 Ma...................................................................................................................174 7-8. Distance in slip direction vs. paleod epth diagram showing geometries for the detachment at the onset of extension at ~53 Ma....................................................175 7-9. Distance in slip directio n vs. paleodepth diagram showing magnitude of offset on the detachment with vari able slip rate from ~53-40 Ma with a geothermal gradient of 35C/km...............................................................................................181 7-10. Geologic map of the ACC showing stru ctural pinning points used to constrain the maximum amount of displacement faci litated by the Anaconda detachment..182 D-1. Major element Harker variation di agrams for samples Ug-1, WG04-114, and DF02-114...............................................................................................................236 D-2. REE trend and trace element spider diagrams for samples Ug-1, WG04-114, and DF02-114...............................................................................................................239

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xiii OBJECT LIST Object page F-1. Geologic map of the Anaconda metamorphic core complex lower plate (AMCCfinalmap.pdf, 22 mb).................................................................................246

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xiv Abstract of Thesis Presen ted to the Graduate School of the University of Florida in Partial Fulfillment of the Requirements for the Degree of Master of Science EXHUMATION AND COOLING HIST ORY OF THE MIDDLE EOCENE ANACONDA METAMORPHIC CORE COMPLEX, WESTERN MONTANA By Warren Calhoun Grice, Jr. August 2006 Chair: David A. Foster Major: Department: Geological Sciences New 40Ar/39Ar thermochronology, U-Pb geoc hronology, and thermobarometry define the tectonic exhumation and cooli ng history of the Middle Eocene Anaconda metamorphic core complex (ACC) of western Montana. Mica 40Ar/39Ar cooling ages obtained from the ACC lower plate (footwall): (1) constrain the age of the onset of extension in the ACC to ~53 Ma; (2) constrai n the duration of extens ion to at least ~5339 Ma; (3) define a lateral cooling age gradie nt, where mica cooling ages decrease to the ESE across the ACC lower plate confirming t op-to-the-ESE directed unroofing of the lower plate; (4) constrain the original geometry of the Anaconda detachment at ~53 Ma to a listric-shaped normal fault comprised of a steep portion (~54-70 ) in the upper brittle crust and a sub-horizontal porti on (~7-12) in the middle crust; and (5) constrain the rate of slip on the detachment to ~0.9 km/Myr during the time interval of ~53-39 Ma. Reconstruction of similar Late Cretaceous granodiorite plutons in the detached ACC upper plate (hanging wall) and lower plat e indicate ~25-28 km of (horizontal)

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xv displacement occurred on the detachment during the Eocene. Thermobarometry and UPb zircon geochronology from the Lake of th e Isle shear zone (LISZ, named herein), structurally beneath the detachment, s how that uppermost-amphibolite facies metamorphism (at ~3.2-5.3 kbar and ~750-850C) ended in the Late Cretaceous (at ~7574 Ma). Pressure constrains from the th ermobarometry indicate the lower plate was exhumed from a maximum crus tal depth of ~10-16 km. The results from this study show the e xhumation and cooling histories of the ACC and the Bitterroot metamorphic core comple x (BCC), located ~70 km west of the ACC, are remarkably similar. Therefore, th e ACC and BCC represent one continuous and integrated extensional system that accommodated large-scale extension in easternmost Idaho and western Montana during the Eocen e. Extension in the ACC and BCC was linked to regional dextral tran stension along the Lewis and Clark Line (LCL), a major strike-slip fault zone located to the nort h of these metamorphic core complexes. Extension in the ACC and BCC, combined with extension in several other metamorphic core complexes north of the LCL, is responsib le for the initial colla pse of the previously thickened Sevier hinterland beginning in the early to middle Eocene immediately following (~1-3 Ma) the end of crustal shorteni ng in the foreland fold-and-thrust-belt to the east. Regional dextral transtension, la rge-scale extension, and exhumation of metamorphic core complexes in the northwe stern United States and southern British Columbia were likely driven by traction cause d by increased obliquity of convergence at the western margin of the North American Plat e beginning in the early to middle Eocene.

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1 CHAPTER 1 INTRODUCTION Metamorphic core complexes are regional extensional structures characterized by large-scale crustal extension and tectonic exhumation of mid-crustal metamorphic and plutonic rocks along low-angle normal brittle-d uctile detachment fault zones (e.g., Davis and Coney, 1979; Coney, 1980). More than th irty metamorphic core complexes have now been documented throughout the North American Cordillera and comprise a relatively narrow and sinuous belt stretchi ng from southern British Columbia into northern Mexico (Coney, 1980; Armstrong a nd Ward, 1991). These metamorphic core complexes formed during early to middle Ter tiary time following a long period of crustal accretion, shortening, and thickening caused by convergence along the western margin of North American during the Mesozoic (Burchfi el et al., 1992; Wernicke, 1992). With the exception of the southernmost metamorphic core complexes, all the core complexes formed in the previously thickened Mesozoic Sevier hinterland, west of the Sevier thinskinned foreland fold-and-thrust belt and Laramide-style basement-involved foreland uplifts (Fig. 1-1; Coney and Harms, 1984; Wernicke, 1992). Large-scale extension in metamorphic core complexes occurred be fore widespread Basin-and-Range-style extension affected the North American Cordil lera (Liu, 2001). Therefore, metamorphic core complexes are responsible for the initial extensional collapse of the previously thickened Cordilleran crust during the Tert iary (Coney and Harms, 1984; Wernicke, 1992; Foster et al., 2001; Vanderhaeghe et al., 2003).

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2 Figure 1-1. Regional geologic map of the No rth American Cordillera. Emphasis is placed on major igneous and tectonic f eatures. Note the location of the Cordilleran metamorphic core complexes located west of the Sevier fold-andthrust-belt. The Anaconda metamorphic core complex (ACC) is located south of the Lewis and Clark line in western Montana. Note the Lewis and Clark Fault Zone = the Lewis and Clark Line. (Modified from Coney, 1980, and Foster et al., 2006a).

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3 Large-scale extension and exhumation of mid-crustal metamorphic and plutonic rocks in metamorphic core complexes in the northwestern United States was linked to regional dextral transtension along the Lewis and Clark line (LCL) during the Eocene (Doughty and Sheriff, 1992; Yin and Oertel, 1995; Foster et al., 2006a). The LCL is a northwest-west-trending ~40-80 km wide zone of steeply-dipping strike-slip, obliqueslip, and dip-slip faults that reaches fr om northeastern Washington to west-central Montana. Many faults of the LCL accommoda ted repeated displacement during multiple tectonic reactivations since th e inception of the line in th e Mesoproterozoic or earlier (Wallace et al., 1990; Sears and Hendrix, 2004) During the Late Cretaceous to early Eocene (~88-55 Ma) the LCL accommodated sinistral displacement and regional transpression related to crustal shortening a nd thickening in the Sevier fold-and-thrust belt of southern Alberta and western Montana. Beginning in the early Eocene (~54-52 Ma), the LCL was reactivated as a regional dext ral transfer structure related to large-scale extension and exhumation of several meta morphic core complexes in the Sevier hinterland, west of the fold-and-thrust belt (Fig. 1-1 and 1-2, Sears and Hendrix, 2004; Foster et al., 2006a). South of the LCL, in eastern Idaho a nd western Montana, large-scale Eocene extension may have been accommodated by exhumation of both the Bitterroot and Anaconda metamorphic core complexes (Fig. 1-2, Foster et al., 2006a). The timing and kinematics of Eocene extension in the Bitte rroot metamorphic core complex (BCC), the more westerly of the two core complexe s, is n well constrained (e.g., Hyndman, 1980; House and Applegate, 1993; House and Hodge s, 1994; Foster and Fanning, 1997; Foster, 2000; Foster et al., 2001; House et al., 2002; Foster and Raza, 2002). Extension in the

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4 BCC began at ~53 Ma and continued to ~25 Ma (based on a large U-Pb geochronology, 40Ar/39Ar, and fission track thermochronological da ta set from the exhumed footwall, Figure 1-2. Tectonic map of nor thwestern United States a nd southern Canada showing major structures related to Tertiary extension. The Bitterroot and Anaconda metamorphic core complexes are located south of the Lewis and Clark Line (LCL). The LCL was active as a regiona l dextral transfer structure between extension north and south of the line during the Eocen e (Modified from Foster et al., 2006a).

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5 Foster et al., 2001; Foster and Raza, 2002; Fo ster et al., 2006a). During this time, the BCC lower plate was exhumed from ben eath a single east-dipping upper-amphibolite facies mylonitic shear zone and overprinting brittle normal detachment fault system by top-to-the-east-southeast directed (110-100) detach ment of the hanging wall (constrained by a large thermochronology data set from the footwall and kinematic indicators and mineral stretchi ng lineations from the mylonite s, Foster, 2000; Foster, et al., 2001; House et al., 2002; Foster and R aza, 2002). Early extension in the BCC was accompanied by upper-amphibolite facies metamorphism at ~6-8 kbar and ~650-750C and localized decompressional anatexis (par tial melting) based on thermobarometry and U-Pb geochronology from migmatites directly beneath the mylonitic shear zone in the eastern parts of the footwall (House et al., 1997; Foster et al., 2001; Foster and Raza, 2002). The thermobarometric data show th at the eastern BCC was exhumed from lower to mid-crustal depths of ~20-25 km (Foster and Raza, 2002; Foster et al., 2006a). Total (horizontal) displacement on th e detachment in the BCC is estimated to be ~40-50 km, based on reconstruction of Cretaceous dioritic plutons in the detached hanging wall with similar dioritic plutons in wester n footwall (Foster et al., 2006a). Extension in the Anaconda metamorphic core complex (ACC, ONeill, et al., 2002; Kalakay and Lonn, 2002; Kalakay et al., 2003; ONeill et al., 2004), located ~70-80 km east of the BCC, is not well constrained, but apparently remarkably similar (Fig. 1-2). Limited U-Pb geochronological and thermoch ronological data suggest extension in the ACC began at ~53 Ma, coincident with the ons et of extension in the BCC, and continued until at least ~47 Ma (ONeill et al., 2004; Fost er et al., 2006a). As in the BCC, the ACC lower plate (footwall) appears to have been exhumed along a single east-dipping

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6 mylonitic shear zone and overprinting brittle detachment fault system which now bounds its eastern side (Kalakay et al., 2003; O Neill et al., 2004). Kinematic indicators and mineral stretching lineations from the mylon ites suggest top-to-the-east-southeast (102108) directed detachment of the ACC uppe r plate (hanging wall) exhumed the ACC lower plate (Kalakay et al., 2003). However, a rather large thermochronological data set is needed from across the ACC to definitely prove east-southeast directed detachment of the upper along an east-dipping detachment system (e.g., Foster and John, 1999). Metamorphosed Lower Belt Supergroup pelitic strata exposed directly beneath the mylonitic shear zone and brittle detachment system in the ACC are migmatitic and show evidence for upper-amphibolite facies me tamorphism and anatexis during high temperature ductile deformation (T. Kalaka y, per comm.). However, the metamorphic grade, age, and relationship of these migma tites to Eocene extension in the ACC have not been established. Geochronological and thermobarometric data are needed from the lower plate migmatites. Within these data, the amount of exhumation in the ACC can not be quantified. In addition, the total amount of displacement facilitated by the ACC detachment during the Eocene has not been established. The apparent similarities in the timing and kinematics of Eocene extension in the BCC and ACC has led some workers to propose the two metamorphic core complexes were part of a continuous integrated exte nsional system that together accommodated large-scale extension south of the LCL dur ing the Eocene (Doughty and Sheriff, 1992; ONeill et al., 2004, Foster et al., 2006a). Following this idea, the BCC and ACC would have been exhumed along separate synthetic (parallel) east-dipping mylonitic shear zones and brittle detachment systems that would have been active at the exact same time. In

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7 this tectonic model a nested more sha llow position for ACC mylonitic shear zone and brittle detachment is proposed (D. Foster, per. comm.; ONeill et al., 2004, their Fig. 8). The primary aim of this study is to test the idea that the BCC and ACC metamorphic core complexes represent a single integrated extensional system responsible for large-scale Eocene extension south of the LC L as has been suggested. In order to test the proposed tectonic model, Eocene extension in the ACC must be constrained. In this study a combination of 40Ar/39Ar thermochronology, U-Pb geochronology, thermobarometry, major-trace element geoche mistry, and field mapping are used to provide a detailed exhumation and cooling hist ory for the ACC. In particular, this study will: (1) confirm the age of the onset extens ion in the ACC; (2) provide constraints on the duration of extension in the ACC (i.e., di d extension in the ACC continue after ~47 Ma?); (3) constrain the Late Cretaceous to Late Eocene cooling history for the ACC lower plate; (4) confirm top-to-the-east-southeast directed tectonic unroofing (detachment) of the ACC upper plate as in the BCC; (5) constrain on the original geometry of ACC detachment to confirm an or iginal east-dipping detachment synthetic to the BCC detachment; (6) constrain the slip rate and magnitude of slip on the ACC detachment system; (7) constrain metamorphi c pressure and temperature conditions and age of upper-amphibolite facies metamorphism in the ACC lower plate; these metamorphic pressure constraints are used to (8) constrain the maximum amount of exhumation in the ACC. The metamorphic pr essure constraints are also used to (9) determine the maximum depth of the ACC det achment in the Eocene to test its proposed nested geometry relative to the BCC detachment.

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8 If the proposed tectonic m odel for large-scale Eocene extension in the BCC and ACC is substantiated then the structural st yle of extension south of the LCL differed significantly from Eocene extension in a nd north of the LCL in northern Idaho, northeastern Washington, and southern British Columbia; as all of these core complexes were exhumed by paired east and west-dipping detachment systems (e.g., the Shuswap metamorphic core complex), not by single as ymmetric detachments (e.g., the BCC). An improved and more complete understanding of Eocene extension in the northwestern Unites States, south of the LCL, is a critical contribution for defini ng the role of plate margin forces, mantle upwelling, and extensi onal (orogenic) collaps e which apparently caused the destruction of an Andean-style orogen that existed in the North American Cordillera from Cretaceous to early Eocene ti me (e.g., Coney and Harms, 1984, Foster et al., 2001; Foster et al., 2006a).

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9 CHAPTER 2 REGIONAL GEOLOGIC BACKGROUND Pre-Mesozoic History Archean ( 2.5 Ga) and Paleoproterozoic base ment underlie western Montana (Foster et al., 2006b). In northwestern M ontana the basement is comprised of the Medicine Hat Block of the southern Archean Hearne Province, which extends beneath southernmost British Columbia, Alberta and Saskatchewan. To the south, the underlying basement consists of the Archean Wyoming Province. These two Archean Provinces are separated by the linear no rtheast-trending Great Falls tectonic zone (GFTZ) which stretches from eastern Idaho to the northeas t into Saskatchewan (ONeill and Lopez, 1985). The GFTZ is mostly Paleoproterozoic in age and has been interpreted to represent the final suturing of the Archean Wyoming Pr ovince and Medicine Hat Block at ca. ~1.91.8 Ga (based on U-Pb zircon crystallizati on ages, Mueller et al., 2003). The WyomingMedicine Hat suture (the GFTZ), along with several other north to northeast-trending Paleoproterozoic collisional belts within the Canadian shield (e.g., the Trans-Hudson orogen) were responsible for the assemblage of the Laurentia cratonic core from ~2.0-1.8 Ga (Hoffman, 1988). The Mesoproterozoic Belt Basin is superi mposed of the western edge of the Archean-Paleoproterozoic basement in wester n Montana and also extends into northern Idaho, northeastern Washington, and southern Alberta (Fig. 2-1). The Belt Basin contains a thick (up to ~18 km) sequence of fi ne-grained quartzites, red to green-colored argillites, argilleous carbonates, and dark-g rey well laminated argi llites and quartzites

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10 Figure 2-1. Regional geologic map of the major Archean basement provinces and Proterozoic suture zones of the Nort h American Cordillera in the western United States and southern Canada. The map also shows the position of the Belt-Purcell Basin superimposed (dark da shed outline) on the western edge of the Precambrian basement provinces in northwestern Washington, northern Idaho, western Montana, and southern Canada. Darkened shapes represent subaerial exposes of Precambrian basement. The Sri = 0.706 line represents the western edge of Precambrian ba sement (from Foster et al., 2006). (Winston and Link, 1993). In the northwestern United States this immense sequence of sedimentary strata is referre d to as the Belt Supergroup a nd in Canada as the Purcell Supergroup. The timing of deposition of the Be lt Supergroup sedimentary strata is now well constrained to the interval of ~1.471.40 Ga by U-Pb zircon crystallization ages from intercalated syn-depositional mafic sills and volcanic tuffs, a large detrital zircon UPb age dataset, and detrital muscovite 40Ar/39Ar cooling ages (Ross and Villeneuve, 2003; and references therein). The origin and Mesoproterozoic tectoni c setting of the Belt

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11 Basin is still controversial (see Winston a nd Link, 1993 for review). There is increasing evidence and a general consen sus among many workers for a extensional rift-related intracratonic basin setting for the Belt Basin, wh ere the subsiding basin was in part, if not fully isolated from an adjacent seaway to the western or northwest (e.g., Winston, 1986; Winston and Link, 1993; Lyons et al., 2000; Ross and Villeneuve, 2003). In western-central Montana the Belt Ba sin narrows significantly to form a prominent east-west-trending structural embaym ent referred to as the Helena embayment (Fig. 2-1 and 2-2). The Helena embayment formed when several major syn-depositional steep normal faults down-dropped the embayment separate from the rest of the Belt basin to the west during the Mesoproterozoic (W inston and Link, 1993). The southern margin of the Helena embayment was bound and dow n-dropped by the east-west-trending Perry Line/Willow Creek fault zone. As a result, the Dillon block of the Archean Wyoming Province was first uplifted along the up-thrown (southern) side of the fault zone (Winston and Link, 1993). The northern margin of the Helena embayment was bound and downdropped along a series of high-angle northwest -west-trending faults which comprised the eastern continuation of the proto-Lewis and Clark line (see below, e.g., Wallace et al., 1990; Sears and Hendrix, 2004). Other major fa ults also active along the eastern segment of the Lewis and Clark line during the Mesopr oterozoic may have included the inferred east-west-trending Jocko and Garn et lines and the Volcanic Va lley fault zone (Fig. 2-2, Winston and Link, 1993). The geometry of the Helena embayment, as defined by the Early bounding fault system, had a major infl uence on subsequent tectonism in western Montana (especially during the Late Mesozoic to Early Cenozoic, Winston et al., 1986; Foster et al., 2001, 2006; Sears and Hendrix, 2004).

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12 Figure 2-2. Simplified tectonic map of Belt Ba sin in the Mesoproterozoic. Map shows the major normal faults active during deposition of the Belt-Purcell Supergroup. The early Lewis and Cl ark likely included the Osburn Fault (OF), the Hope Fault (HF), the Jocko Line (JL), and Garnet Line. The JL and GL are inferred (see Winston and Link, 1993). The Volcano Valley Fault (VVF) and the Willow Creek Fault Zone (WCF) were responsible for downdropping the Helena Embayment duri ng deposition of the Belt-Purcell Supergroup. The 87Sr / 86Sr = 0.706 line from Sears and Hendrix (2004) and represents the western edge of known Precambrian basement. The Great Falls Tectonic Zone (GFT) was not active dur ing deposition of the Belt-Purcell Supergroup Mesoproterozoic (Modifie d from Winston and Link, 1993). Following the deposition of the Belt S upergroup sedimentary strata, during the Late Neoproterozoic (ca. ~800-700 Ma), a majo r rift zone developed along the entire western margin of Laurentia as the proposed pre-Gondwanan supercontinent Rodinia began to break apart (e.g., Meer t and Torsvik, 2003). As a resu lt, the western side of the Belt Basin was truncated as the adjacent landmass rifted away (the identity of the western landmass is highly controversial; see Karlst rom et al., 2001 for review). As rifting continued, a paleo-Pacific Ocean opened and a passive margin sequence was deposited beginning with the deposition of the Late Neoproterozoic to Early Cambrian-aged

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13 Windermere deep-water sequence along the en tire western margin of Laurentia but mainly outboard (west) of the Belt Basin in western Montana (Bur chfiel et al., 1992; Winston and Link, 1993). During the Middle Cambrian, a thick mi ogeoclinal sequence (continental shelf deposits) was deposited over the eastern Winde rmere sequence and further inboard onto the craton during a major transgression (Burchfi el et al., 1992). In western Montana, the earliest phase of this transg ression is represented by the Middle Cambrian Flathead Sandstone which lies unconformably over uppe r Belt Supergroup stra ta in much of western Montana (Emmons and Calk ins, 1913; Winston and Link, 1993). Mesozoic History Beginning in the Early Triassic, an activ e convergent margin developed along the western margin of Laurentia (B urchfiel et al., 1992). With continued convergence, from the Early Triassic to Early Cretaceous, severa l exotic terranes (including a number of volcanic arc complexes) were accreted to the western edge of Laurentia (e.g., the Wrangellia terrane and the Wallowa-Seven Devils terrane adjacent to the Western Idaho Suture, Hamilton, 1978). Despite the major c ontractional deformation associated with convergence at this time, the Belt Basin a nd its Paleozoic miogeoclinal sedimentary cover sequence remained largely stable until the Late Early to Late Cretaceous (Burchfiel et al., 1992). By the Late Early to Late Cretaceous, cr ustal shortening and thickening related to the ongoing convergence along the western margin of Laurentia migrated eastward into present-day northeastern Idaho and western Montana. Thru st faulting began first in eastern Idaho at ~105 Ma and moved progr essively eastward into western Montana before ending at ~55 Ma (Hyndman, et al ., 1988; Constenius, 1996; Sears and Hendrix,

PAGE 29

14 2004). In western Montana, the east-direc ted crustal shortening and thickening was facilitated predominately by low-angle thinskinned basement-detached Sevier-style thrust faulting superimposed on the Mesopr oterozoic Belt Basin (Allmendinger, 1992; Winston and Link, 1993); this deformation was pa rt of the Sevier orogen which stretches from British Columbia into southern Nevada (Burchfiel et al., 1992). As a result, many thin-skinned thrusts in western Montana invol ve thick sequences of Belt Supergroup and Paleozoic sedimentary strata thrust eastwa rd over younger Belt, Paleozoic, and Mesozoic sedimentary strata (Winston, 1986; Foster et al., 2001). In addition, because thrusting was superimposed on the Belt Basin, a prominent east-directed structural bulge formed in the thrust belt in the original Helena embaym ent referred to as the Helena thrust salient (Fig. 1-2 and 2-3; Winston, 1986). Crustal Shortening along the Lewis and Cl ark Line during the Late Cretaceous Thin-skinned Sevier-style thrusting in western Montana was accommodated by the eastward displacement of three major thrust slabs (or plates) along the Lewis and Clark Line (LCL): 1) the Lewis-Eldorado-Hoadle y (LEH) slab in northwestern Montana and the 2) Sapphire and 3) Lombard slabs in s outhwestern Montana (Constenius, 1996; Sears and Hendrix, 2004). The LCL, a complex zone of west-to-northw est-trending steeply dipping strike-slip, dip-slip and oblique slip faults, separates the LEH slab in the north from the Sapphire and Lombard slabs to th e south (Fig. 2-3, Wallace et al., 1990). As noted above, some faults within the LCL likel y originated during subsidence of the Belt Basin in the Mesoproterozoic. However, se veral workers show evidence for reactivation of these old faults as well as inception of younger faults along the LCL during the Mesozoic (e.g., Harrison et al., 1972; Hyndma n et al., 1988; Wallace et al., 1990; Stewart and Crowell, 1992; Sears and Hendrix, 2004). Most recently, Sears and

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15 Figure 2-3. Tectonic map of the North Am erican Cordillera of western Montana, northern Idaho, northeastern Washingt on, and southern Canada during the Late Cretaceous. Major thrust slabs in western Montana were comprised of the Sapphire and Lombard slabs south of the Lewis and Clark Line and the Lewis-Eldorado-Hoadley slab to the nor th. Note the eastward budge in the fold-and-thrust directly east of the Sa pphire and Lombard sl abs referred to as the Helena Salient. The ISr = 0.706 line represents the western edge of Precambrian basement (Modified from Sears and Hendrix, 2004). Hendrix (2004) show that the LCL was reactivated as a major left-Lateral transpressional shear zone from the Late Cretaceous to the Early Eocene (~88-55 Ma). In their model, left-Lateral motion along the LCL accomm odated clock-wise rotation and eastward displacement of the LEH, Sapphire, and Lombar d thrust slabs in western Montana from

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16 Late Cretaceous to Early Eocene time. In a ddition, Sears and Hendrix (2004) attribute development of the Helena and Alberta (s outhern Alberta and nor thwestern Montana) thrust salients to clock-wise rotation of these major thrust slabs along the LCL during this time. Late Cretaceous Magmatism Voluminous magmatism was broadly sync hronous with crustal shortening and thickening in eastern Idaho and western Mont ana. In eastern Idaho and westernmost Montana the Bitterroot lobe (~14,000 km2) of the Idaho-Bitterroot Batholith was emplaced into the accreted Seven Devils-W allow terrane and Belt Supergroup largely from ~120-52 Ma (see Fig. 1-1, Bickford et al., 1981; Hyndman, 1984; Foster and Fanning, 1997; Foster et al., 2001) Early (~120-70 Ma) deep level quartz diorite and tonalite plutons of the western Bitterroot lo be were emplaced at crustal depths up to ~25 km (Hyndman et al., 1988; Foster et al., 2001). In the central and ea stern Bitterroot lobe main phase granodiorites and two-mica gr anites were emplaced at progressively shallower crustal depths to the east (~2510 km) between 65-52 Ma. In addition, several younger alkali-feldspar granite plutons were em placed in the older batholith intrusions from ~50-46 Ma at very shallow crustal de pths ranging from ~7-1.5 km (Foster et al., 2001). East of the Bitterroot lobe of the Idaho-Batholith Batholith several smaller batholiths, stocks, and plutons were emplaced into the Sapphire and Lombard thrust slabs from ~80-65 Ma at relatively shallow crus tal depths (~1-15 km, Hyndman et al., 1988; Kalakay et al., 2001). Major intrusions in th is region include the Sapphire Batholith, the Chief Joseph Batholith, the Flint Creek plutons and the Boulder Bat holith (Fig. 2-4).

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17 The Boulder Batholith, the largest of these intrusions (~4,000 km2), intruded its own volcanic cover (the El khorn Mountains Volcanics) indicating a very shallow emplacement depth, probably on the order of ~1-10 km (Tilling et al., 1968). Individual plutons and stocks within thes e large intrusions often exhibi t a close spatial and temporal relationship with thrust faulting in the Sapphi re and Lombard thrust slabs. In several places within the Flint Creek and Anaconda-Pintl ar Ranges (i.e., the Sapphire thrust slab) Late-Cretaceous plutons were preferentia lly emplaced along thrust faults ramps (Hyndman et al., 1975; Kalakay et al., 2001). Older plutons commonly exhibit strong solid-state deformation while slightly younge r plutons emplaced along the thrusts (some cross-cutting the thrusts) are undeformed; th is relationship indicates the older plutons were probably emplaced during thrusting while emplacement of the younger plutons post-dated the thrusting (e.g., Hyndman et al 1975; Hawley, 1975; Wallace et al., 1992; Kalakay et al., 2001). Kalakay et al. (2001) reports similar relationships between LateCretaceous plutons and thrust faulting in easte rn Pioneer Batholith along the eastern front of the Grasshopper thrust slab located directly south of the Sapphire thrust slab in southwestern Montana. Similarly, the Boul der Batholith was emplaced into a major thrust fault ramp system within the Lombard thrust slab in the Helena thrust salient at ~80-75 Ma (Fig 4., based on SHRIMP U-Pb zircon crystallizati on ages, Lund et al., 2002). Lageson et al. (2001) suggest that em placement of the Boulder Batholith (and the Elkhorn Mountains Volcanics) at this time created a supercr itical taper geometry in the Lombard thrust slab facilitating further east-directed displace ment of the slab.

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18 Figure 2-4. Geologic map of wester n Montana, south of the Lewis and Clark line. Emphasis is on the location and names of Late Cretaceous through middle Eocene granitic ba tholiths, stocks, plutons, and volcanic fi elds in the Sapphire thrust plate and the frontal imbricate thrust zone which lies to the east (from Wallace et al., 1992).

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High Grade Metamorphism during the Cretaceous Grover et al. (1992) and H ouse et al. (1997) document three major metamorphic events that affected the areas adjacent to th e Bitterroot lobe of the Idaho-Bitterroot Batholith during the Late Early to Late Cretaceous. (1) Regional prograde metamorphism at ~100-80 Ma was synchronous with crustal shor tening and partly coincident with emplacement of deep level pl utons in the western Bitterroot lobe (~12070 Ma, Foster et al., 2001) Pressure and temperat ure conditions during this metamorphism have been estimated (based on phase assemblages and quantitative thermobarometry) to be at Middle-amphi bolite facies conditions, ~5-6 kbar and ~500600C (House et al., 1997 and references therein). (2) Upper-amphibolite facies metamorphism occurred synchronous to emplacement of main phase plutons in the Bitterroot lobe at ~64-53 Ma. Metamor phic grade decreases from upper-amphibolite facies conditions (peak metamorphic conditi ons at ~6-8 kbar and ~650-750C) directly adjacent to the batholith to lower greenschis t facies conditions at a distance of ~30 km (House et al., 1997; Foster et al., 2001) Local anatexis accompanied the upperamphibolite facies metamorphism near the batholith. Anatexis of Belt-equivalent quartztwo feldspar gneisses was fac ilitated by the breakdown (dehydr ation) of muscovite within underlying semi-pelitic schist under sillimanite -zone conditions (Foster et al., 2001). (3) Later relatively low pressure (~4-6 kbar) high temperature upper-amphibolite facies metamorphism and anatexis during isothermal decompressional accompanied exhumation of the Bitterroot metamorphic core complex at ~53-48 Ma (see below, House et al., 1997; Foster et al., 2001). Quantitative thermobarometric data were not available from areas east of the Bitterroot-Idaho Batholith prior to this st udy. Some workers provide broad pressure-

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20 temperature estimates based on the metamor phic phase assemblages of metasedimentary strata adjacent to plutons. These estimates indicate localized low pressure Middle to upper-amphibolite facies metamorphism at ~1-4 kbar and ~500-700C accompanied emplacement of the plutons during the La te Cretaceous (~80-70 Ma, Hyndman, 1988, see previous pressure-temperature constrai nts summarized in a Later section). Eocene History An episode of large-scale extension began in the Ea rly to Middle Eocene with inception of numerous metamorphic core complexes along the previously thickened Sevier hinterland, north of the Snake River Pl ain (SRP) in present-day southern British Columbia, northeastern Washington, northern Idaho, and western Montana (Foster et al., 2001, Foster et al., 2006a). The large-scale co re-complex related extension immediately followed the end of crustal shortening (ca. ~55 Ma) in the fold-and-thrust belt located directly to the east of the hinterland (Fig. 1-2, Constenius 1996; Foster et al., 2001; Sears and Hendrix, 2004). Large-scale extension in these metamorphic core complexes was facilitated by displacement al ong low-angle brittle/ ductile normal-sense detachment fault zones comprised of a brittle upper crustal port ion rooted to a mid-crustal ductile portion; the mid-crustal portion of these detachment s are characterized by mylonitic shear zones (Coney, 1980; Lister and Davis, 1989; We rnicke, 1992; Foster et al., 2001). Early extension in the metamorphic core complexes of southern British Columbia and the northwestern United States was accompanied by the voluminous KamloopsColville-Challis-Absaroka (KCCA) magmatism in the ba ck-arc region, ~500-1000 km inboard of the convergent plate margin (A rmstrong and Ward, 1991; Morris et al., 2000; Breitsprecher et al., 2003). The (KCCA) magmatism is expressed by a belt of calcalkaline and alkaline volcanic rocks exte nding from British Columbia into northwestern

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21 Wyoming and roughly overlaps with areas aff ected by extension in the metamorphic core complexes. The exhumed lower plates of many of the Eocene metamorphic core complexes in southern British Columbia and the northwestern United States are characterized by voluminous granitic plutons the same age as the (KCCA) volcanic rocks (Armstrong and Ward, 1991; Foster et al., 2006a). Regional Transtension and Large-scale Crustal Extension along the Lewis and Clark Line Large-scale crustal extension in the meta morphic core complexes of northeastern Washington, northern Idaho, a nd western Montana was linked kinematically to regional dextral transtensional along the Lewis a nd Clark line during the Eocene (LCL, Doughty and Sheriff, 1992; Yin and Oertel, 1995; Foster et al., 2006a). The LCL is comprised of a northwest-west trending ~40-80 km wide zone of steeply-dipping strike-slip, obliqueslip, and dip-slip faults that stretches from northeastern Washington into west-central Montana (a distance greater than 800 km). Several major faults along the LCL have accommodated repeated displacement during tect onic reactivations since inception of the line during the Mesoproterozoic or earlier (Wallace et al., 1990; Sears and Hendrix, 2004). Regional dextral transtension and deve lopment of metamorphic core complexes along the LCL began in Early to Middle Eocen e time (at. ~54-52 Ma) after a long-lived episode of sinistral transpressional during th e Late Cretaceous to early Eocene (ending at ~55 Ma) associated with east-directed crusta l shortening in the fold-and-thrust-belt of western Montana (Foster et al ., 2001; Sears and Hendrix, 2004; Foster et al., 2006a). The detachment fault zone of the metamorphic co re complexes along the LCL are structurally linked to one another via strike -slip and/or oblique-slip splays faults of the LCL; these

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22 splay faults exhibit major Eocene aged dextra l offset (Fig. 1-2, Hendrix and Sears, 2004; Foster et al., 2006a). Although the timing of large-scale Eocen e extension in metamorphic core complexes along the LCL is very similar, th ere are some significant differences in the kinematics of extension in these metamorphic core complexes north and south of the LCL. To the north, metamorphic core complexes were exhumed by paired (or symmetric) east and west dipping detachment fault zone by ENE-WSW directed tectonic unroofing (Fig. 1-2; e.g., Okanagan, Kettle and the Priest River metamorphic complexes). To the south, the BCC was exhumed by a single (or asymmetric) east dipping detachment fault zone by ESE directed unroofing (Foster et al., 2006a). It has been proposed that the ACC was exhumed along east-dipping detachment fault zone by ESE directed unroofing as well (Kalakay et al ., 2003; ONeill et al., 2004, Foster et al., 2006a). The Clearwater metamorphic core co mplexes is located within the LCL and occupies a relay between tw o major strike-slip faults which link the Priest River metamorphic core complex and the BCC (Fig. 1-2). The Clearwater metamorphic core complexes was exhumed along single east-dippi ng detachment fault zone by top-to-theeast directed unroofing, similar to the BCC (Foster et al., 2006a). .

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23 CHAPTER 3 THE ANACONDA METAMORP HIC CORE COMPLEX Structural-metamorphic Domains of the Anaconda Metamorphic Core Complex: The Anaconda metamorphic core complex (ACC) is located within the collapsed Sevier hinterland province in western Montana, south of the Lewis and Clark line, east of the Bitterroot metamorphic core complex, and west of the Boulder Batholith, and Helena salient of the fold-and-thrust belt (Fig. 1-1 and 1-2). As in other Cordilleran metamorphic core complexes (e.g., Coney, 1980), the ACC is subdivided into three structural-metamorphic domains: (1) the high grade metamorphic and plutonic lower plate (footwall), (2) the unmetamorphosed or low-grade metamorphic, brittle faulted upper plate (hanging wall), and (3) the det achment fault zone which juxtaposes and separates the upper and lower plates. In this chapter general descriptions are provided for each of the three structuralmetamorphic domains of the ACC. Detailed de scriptions are provided for the Lake of the Isle shear zone, a km-scale upper-amphibolite facies ductile shear zone mapped in the northeastern Anaconda-Pintlar Ra nge during this study. Deta iled descriptions of the individually mapped ACC lower plate units and a new 1:24,000 scal e geologic map of the study area are provided in A ppendix B and D, respectively. The Lower Plate Rocks of the ACC lower plate are partly exposed in the cores of two broad antiformal flexures, or domes, in the Flin t Creek and northern Anaconda-Pintlar Ranges that lie structurally beneath the detachme nt. These two antiforms are separated by a

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24 synformal trough west of Anaconda, MT where upper plate rocks partly cover the lower plate rocks (Fig. 3-1, ONeill et al., 2004). Lower plate rocks in the Flint Creek and Anaconda-Pintlar Ranges are comprised part ly of metamorphic equivalents to the Mesoproterozoic Belt Supergroup, Middle Camb rian, and in some areas, to Devonian to Lower Cretaceous sedimentary strata. All of these metasedimentary strata were intruded by multiple generations of Late Cretaceous to early-middle Eocene batholiths, plutons, stocks, dikes, and sills (Fig. 3-1 and 32, Emmons and Calkins, 1913; Csejtey, 1963; Stuart, 1966; Desmarais, 1983; Heise, 1983; Wallace et al., 1992; Lonn et al., 2003). A new 1:24000 scale geologic map was produced fr om the ACC lower plate located in the field area of this study. The map and a deta iled description of th e mapped lower plate units are included in Appendix F and B, respectively. In the Flint Creek Range, the intrusive ro cks comprise epizona l biotite-hornblende granodiorite and granitic plutons of the Late Cretaceous Mount Powell and Philipsburg Batholiths and the Royal Stock which intrude mostly Cambrian and Lower Cretaceousequivalent metasedimentary st rata and Belt-equivalent meta sedimentary strata in a few areas (Fig. 3-1, Allen, 1966; St uart, 1966; Hyndman et al., 1972; Lonn et al., 2003). In the Anaconda-Pintlar Range, both Late Cretaceo us and early to middle Eocene intrusive rocks intrude mostly Belt and middle Camb rian-equivalent metasedimentary strata (Desmarais, 1983; Wallace, et al., 1992; Lonn et al., 2003). The early to middle Eocene intrusive rocks are much more voluminous and form a southward widening belt of biotite granodiorite, biotite granite, and two-mica granite batholith s, stocks, and plutons along the entire length of the An aconda-Pintlar Range (Fig. 3-2, Desmarais, 1983; Wallace et al., 1992).

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25 Figure 3-1. Geological Map of the Anaconda metamorphic core complex (ACC) in western Montana. The exhumed ACC lo wer plate (footwall) highlighted in red is exposed in the Anaconda-Pintlar eastern Flint Creek Ranges. The detachment zone bounds the eastern side of the lower plate and consists of the Eocene greenschist facies mylonite and overprinting brittle detachment. The ACC upper plate (hanging wa ll) lies mostly east of the lower plate and detachment zone with limited exposes ju st west of Anaconda, MT between the Anaconda-Pintlar and Flint Creek Ranges. RS = Royal Stock, Pb = Philipsburg Batholith, MPb = Mt. Powell Batholith, SLP = Storm Lake Stock pluton, SP = Sapphire Batholith, CJ = Ch ief Joseph Batholith, PB = Pioneer Batholith, BB = Boulder Ba tholith, LCV = Lowland Creek Volcanic Field, EHV = Elkhorn Mountains Vo lcanic Field. The curre nt study area is marked by the dashed box (Modified from Foster et al., 2006a).

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26 Figure 3-2. Geologic cross section of the An aconda metamorphic core complex, western Mont ana. The Eocene mylonite and brittle detachment dip gently to the east beneath the Deerlodge Vall ey and project beneath the Boul der Batholith as indicated by Amoco exploration wells and seismic refl ection data (see text for de tails). Equal-area stereonet projection shows the orientation of mylonite stretchi ng lineations (solid squares) an d fault slickenlines (open circles) from the detachment zone (Stereonet data and cross secti on from courtesy of Kalakay).

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27 Major structures exposed with the ACC lower plate include numerous folds and low-angle thrust faults relate d to crustal shortening along th e eastern edge of the Sapphire (Skalkaho) thrust plate during the Late Cretaceous (Hyndman, 1975; Hyndman et al., 1988; Wallace et al., 1992; Sears and Hendrix, 2004). The Georgetown thrust (Emmons and Calkins, 1913), the largest of the thru st fault structures, places metamorphosed Helena Formation (middle Belt Carbonate -equivalent) over Devonian, Pennsylvanian, and Mississippian metasedimentary strata (>700 0 m of stratigraphic offset) and marks the western margin of the ACC lower plate from central Anaconda-Pintlar Range to the northern Flint Creek Range (Fig. 3-2, Wall ace et al., 1992; ONeill et al., 2004). The northern end of this thrust is cut by the Philipsburg Batholith in the western Flint Creek Range which yielded concordant hornblende and biotite K-Ar cooling ages of ~77-72 Ma (Hyndman et al., 1972). Several other thrust faults and related folds involving Beltequivalent through Cretaceous metasediment ary strata have been documented in throughout the Flint Creek and Anaconda-Pintl ar Ranges (e.g., Emmons and Calkins, 1913; McGill, 1959; Flood, 1974; Heise, 1983 ; Baken, 1984; Wallace et al., 1992; Lonn et al., 2003). A common characteristic in both the Anaconda-Pin tlar and Flint Creek Ranges is the close spatial and temporal re lationship between thrusting and emplacement of Late Cretaceous plutons (e.g., the Philip sburg Batholith and Georgetown thrust, Emmons and Calkins, 1913; Wallace et al., 1992). Other structures found within the ACC lowe r include high-angle, sometimes listricshaped, normal faults associated with th e brittle detachment found mostly along the eastern flanks of the Flint Creek and Anac onda-Pintlar Ranges (see below). The largest normal fault exposed in the ACC lower plate is the Hidden Lake-Dry Creek fault zone

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28 (HLCDFZ), a large eastdipping, listric-shaped, locally du ctile normal fault zone that strikes parallel to and lies di rectly east of the Georgetown thrust (Fig. 3-1, 3-2). The HLDCF shows >450 m of offset in some areas and may have been active in the LateCretaceous during intrusion of the Dora Thorn pluton of the Philipsburg Batholith (Buckley, 1990) and during the Eocene, as list ric splays of the HLDCF form ductile shear zones in the middle Eocene Pintlar Creek Bat holith in the northern Anaconda-Pintlar Range (Wallace et al., 1992). The shallow portion of these splay faults may have been part of the original east-d ipping detachment breakaway z one or an east-dipping brittle normal fault synthetic to the basal de tachment (ONeill et al., 2004). The Mesoproterozoic Belt through Cretace ous-equivalent metasedimentary strata exposed in the ACC lower plate have been s ubjected to at least two major episodes of high grade metamorphism: (1) pervasive regional metamorphism that predated the emplacement of voluminous epizonal Late Cret aceous to Middle Eocene intrusions and (2) high temperature, lower pressure cont act metamorphism associated with the emplacement of the Late Cretaceous to Middle Eocene (Emmons and Calkins, 1913; Csejtey, 1963; Stuart, 1966; Desmarais, 1983; Heise, 1983; Hyndman et al., 1988; Wallace et al., 1992; Kalakay et al., 2003; Grice et al., 2005). In the Flint Creek Range, a pervasive regional upper-amphibolite facies co rdierite-bearing assemblage, found mostly within pelitic Cretaceous-equivalent phylli te, is overprinted by andalusite-bearing assemblages in contact aureoles surrounding the Late Cretaceous intrus ions (Stuart, 1966; Buck, 1990). The earlier regional metamorphic assemblages show an overall increase in metamorphic grade from greenschist to upper-am phibolite facies from west to east across the Flint Creek Range (Stuart, 1966). In the Anaconda-Pintlar Range, upper-amphibolite

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29 facies sillimanite-cordierite bearing assemblages are common in Lower Belt-equivalent pelitic metasedimentary strata (Greyson Fm .) surrounding Late Cretaceous granodiorite and quartz diorite plutons and si lls. These metasedimentary strata are locally migmatitic and characterized by granitic leucosome, a bundant sillimanite, cordierite, garnet, Kfeldspar, biotite, and no primary muscovite (Desmarais, 1983; Grice et al., 2005). This phase assemblage overprints an earlier higher pressure kyanite-bearing assemblage in the current study area in the northeastern Anac onda Pintlar Range (see description of the LISZ below, Kalakay et al., 2003; Grice et al., 2005). In additi on, andalusite-bearing assemblages overprint cordierite-sillimanite-bearing assemblages in some pelitic metasedimentary strata adjacent to early to middle Eocene intrusions in the northern Anaconda Pintlar Range (Emmons and Calkins, 1913). Andalusite assemblages were not documented in the contact aureoles of Late Cretaceous intrusions in the current study area. As in the Flint Creek Range, the re gional metamorphic grade shows a general increase from middle to uppermost amphibolite facies from west to east across the Anaconda-Pintlar Range (Kalakay et al., 2003). A summary of previous pressuretemperature estimates from the lower plate, along with new thermobarometry from LISZ in this study, are provided in chapter 5. The Detachment Fault Zone The high grade metamorphic and plut onic rocks of the lower plate and unmetamorphosed upper plate rocks of the ACC are juxtaposed and separated by an eastdipping low-angle brittle detachment that exhi bits significant topto-the-east-southeast displacement (Fig. 3-1 and 3-2, ONeill et al ., 2002; Kalakay et al., 2003; ONeill et al., 2004). The detachment was first recognized and mapped by Emmons and Calkins (1913) who noted its striking similarities with the great Bitterroot fault, which was later

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30 defined to be the low-angle detachment th at bounds the Bitterroot metamorphic core complex to the west (Hyndman, 1980). The detachment now has a documented strike length of >100 km stretching from the norther n flank of the Flint Creek Range south to the Bighole Valley along the eastern flanks of the southern Anaconda-Pintlar Range (Kalakay et al., 2003; ONeill et al., 2004). In the north, the detachment terminates into steeply-dipping strike-slip and oblique-slip splay faults of the Ninemile fault, part of the greater Lewis and Clark line (Foster et al ., 2006a). The southern termination of the brittle-ductile detachment has not been well es tablished. Along the eas tern flanks of the Flint Creek and Anaconda-Pintlar Ranges the detachment dips gently (~10-30) beneath the Deerlodge Valley; the gentle dip of th e detachment is collaborated by industry exploration wells that intersected greenschist mylonites at the base of the Tertiary basin fill in the western Deerlodge Valley at depths of 5 km (Fig 3-2, McLeod, 1987). In addition, the downward projection of the det achment aligns well with sub-horizontal seismic reflectors beneath the Boulder Bathol ith, suggesting the detachment flattens with depth and continues to the eas t (Fig. 3-2, Vejmelek and Sm ithson, 1995; Foster et al., 2006a). The detachment is not well exposed along the western margin of the ACC; however, it is inferred in several places by th e juxtaposition of brittlely faulted upper plate rocks with ductily deformed metamor phic rocks and plutonic rocks. The western part of the detachment probably originated as a series of east-dipping listric-shaped normal faults east of a breakaway zone, which is inferred to be directly east of the Georgetown thrust (ONeill et al., 2004; Foster et al., 2006a). In the northeastern Anaconda-Pintlar Ra nge, within the current study area, the brittle detachment overprints a ~300-500 m thick greenschist facies mylonitic shear zone

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31 comprised mostly of stretched two-mica gr anite, biotite granite, and granodiorite and minor micaeous quartzite (Fig. 3-3, Emmons and Calkins, 1913; Kalakay et al., 2003; Appendix D). Strain is heterogeneous in th e granitoid mylonites and distributed into numerous 1-2 m thick ultramylonite zones se parated by 5-15 m thick zones of mylonite and protomylonite (Foster et al., 2006a). The metamorphic grade of the granitoid mylonites is indicated by br ittle fractured feldspar porphyroclasts surrounded by a matrix of plastically deformed quartz (Kalakay et al., 2003). In addition, micaeous quartzite mylonites exhibit unannealed quartz grains with well-deve loped undulatory extension in thin section (Fig. 3-4a); these features are indicative of lowe r to middle greenschist facies metamorphism at temperatures < 400-450C (Wells et al., 2000). In addition, the greenschist mylonites contain kinematic indicators and shallow plunging mineral stretching lineations that s how top-to-the-eastsoutheast sense of movement; mineral lineations in both the granitoid and micaeous quartzite mylonites are comprised of stretched quartz ribbons (Fig 3-4a and 4b, 102-108, Kalakay et al., 2003; ONeill et al., 2004). The greenschist facies mylonites exposed in the Mill and Clear Creek drainages are cut by series of closely spaced (~0.1-1 km) eas t-dipping listric-shaped normal faults that commonly become sub-horizontal with depth and tangential to th e highly strained ultramylonite zones in the mylonites (Kalakay et al., 2003). In some places, the deeper portions of these brittle faults terminate in to the detachment (ONeill et al., 2004). Good examples of these faults are found on the west ern walls of the Clear Creek drainage and along the continental divide separating the Mill and Tenmile Lake drainages (Fig. 7-5, Appendix D). Slickenline striations on th ese fault surfaces indicate top-to-the-east-

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32 southeast (100-110) slip, consistent with mo tion in the greenschist mylonites (Kalakay et al., 2003; Foster et al., 2006a). Figure 3-3. Photomosaic of the Anaconda meta morphic core complex. Photo directed to the northwest showing the brittle detachment (dashed and barbed line) which juxtaposes unmetamorphosed Tertiary upper plate rocks w ith the Eocene greenschist facies mylonite zone and high grade lower plate rocks. To the north, along the eastern flank of the Flint Creek Range, Eocene mylonites formed in the eastern parts of the Mo unt Powell Batholith and Royal Stock, both comprised mostly of Late Cretaceous gra nodiorite and granite plutons (Allen, 1966; Lonn et al., 2003; ONeill et al., 2004). These myl onites are cut by high-angle normal faults similar to those exposed in the Clear and Mill Creek drainages in the northeastern Anaconda-Pintlar Range (ONeill et al., 2004). To the south, along the eastern flanks of the central and southern Anaconda-Pintlar Range, Wallace et al. (1992) mapped and described low grade mylonites in two mica granites and granodiorites of the Middle Eocene Pintlar Creek Batholith. These mylon ites are cut by a series of northeast-trending

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33 steeply-dipping normal faults related to the de tachment that juxtapose the mylonites with unmetamorphosed Tertiary fluvial deposits. Figure 3-4. Greenschist facies mylonites. A) Outcrop photo of stretched granodiorite with greenschist facies mylonitic fabr ic, upper Clear Creek drainage. Pencil is oriented parallel to the stretchi ng lineation (T/P = 100/10). B) Photomicrograph of mylonitic micaeous qua rtzite, Short Peak, right is to the ESE. Asymmetric mica fish with top-to-the-ESE sense of shear. Note undulatory extinction in the large deform ed quartz grains. Cross polar light; field of view is ~2 mm. The Upper Plate Structurally above the high grade metamo rphic-plutonic lower plate, greenschist facies mylonitic shear zone, and brittle detachment, the eastern ACC upper plate is mostly composed of an array of asy mmetrical fault-bound basins filled with A

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34 Figure 3-5. Outcrop photos of la te listric-shaped brittle norm al faults in the greenschist facies mylonite zone. A) Photo taken from the southeastern flank of Short Peak looking towards the northwest showi ng a listric-shaped brittle fault that juxtaposed mylonitic micaeous quartzite with biotite granodiorite. Symbols: Dot within the circle indicates previo us fault motion out-of-the-page and X within the circle indicates previous fa ult motion into-the-page. B) Photo taken from eastern side of the upper Clear Creek drainage looking south showing an east-dipping listric-shaped br ittle fault cutting greenschist facies two-mica granite mylonite (photo in B is courtesy of David Foster). unmetamorphosed syn-extensional Tertiary sedimentary, volcanic lastic, and volcanic rocks (Fig. 3-1, 3-2). The faults that bound th e basins are listric-shaped and sometimes

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35 sole into the low-angle detachment at depth. In the Deerlodge Valley, the stratagraphically lowest rocks in the basins c onsists of moderately west tilted (~50-60), poorly sorted, and poorly consolidated congl omerates, sandstones, breccias and megabreccias (Kalakay et al., 2003; ONeill et al ., 2004). These strata grade upwards into progressively less tilted (~0-25) volcanic la va flows, volcanic tuffs, and volcaniclastic units that are correlated with the early to middle Eocene Lowland Creek Volcanic (LCV) sequence (~54-48 Ma, based on 40Ar/39Ar cooling ages from the volcanic units by Isopolatov, 1997; Lewis, 1990; Kala kay et al., 2003). The significant upward decrease in tilt of the basin fill strata indicates deposition during brittle extension of the ACC upper plate in the early to middle Eocene (Kalakay et al., 2003). Clasts of greenschist facies two-mica granite and granodiorite mylonite were found in stratagraphically high volcaniclastic lahar deposits in the upper plat e, directly south of the current study area. However, the age or volcano-stratigraphic co rrelation of these volcaniclastic deposits is not yet well established (Kalakay et al., 2003; Foster, per comm.). Description of the Lake of the Isle Shear Zone A sinuous, km-scale ductile shear zone was documented and mapped in the ACC lower plate within the nort heastern Anaconda-Pintlar Ra nge during this study. The ductile shear zone strikes approximately eas t-west and outcrops over a broad area from immediately west of Storm Lake ~15 km to th e east near the base of Mount Haggin in the upper Mill Creek drainage. In the Mount Haggin area, the east ernmost part of the ductile shear zone is truncated by granitoids and greenschist my lonites associated with the Eocene detachment. The ductile shear zone ha s been named the Lake of the Isle shear zone (LISZ) herein after a small lake in th e central part of the shear zone (see Appendix D).

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36 Mesoproterozoic Belt Supergroup and Middle Cambrian-equivalent metasedimentary in the LISZ have been subj ected to pervasive ductile deformation under middle to upper-amphibolite facies conditions and exhibit a well-developed transposed metamorphic foliation throughout the structure; th e dip of the foliation varies from gentle to moderate in east and west and is sub-vert ical in the central LISZ (Appendix F). As a result, the metasedimentary strata deformed in the LISZ are strongly attenuated (ductily thinned) and characterized by dramatically reduced original stra tigraphic thicknesses, common mesoscopic boudins, and recumbent, lo cally isoclinal mesoscopic folds. Late Cretaceous and early to middle Eocene plutons, di kes, and sills intrude the LISZ along its entire length. The spatial re lationship between thes e intrusions and the LISZ is described below along with several aspects of the ductile deformation and high grade metamorphism that were briefly summarized here. Description and Distribution of Me tamorphic and Structural fabrics Metamorphic foliations and gneissic banding Metasedimentary strata deformed and meta morphosed in the LISZ are correlated with metamorphosed Mesoproterozoic Belt Supergroup and Middle Cambrian sedimentary strata (Appendix B). Belt Supe rgroup-equivalent meta sedimentary strata mapped in the LISZ include the metamorphosed Greyson Formation (Lower Belt), Ravalli Group, Helena Formation (Middle Be lt Carbonate), and Missoula Group. Middle Cambrian-equivalent metasedimentary st rata mapped include the metamorphosed Flathead Formation, Silver Hill Formation, and Hasmark Formation. All of the meta-Belt and meta-Cambrian meta sedimentary strata in the LISZ exhibit a metamorphic foliation and/or gneissic (com positional) banding. The pelitic schist correlated with Greyson Formation (Lower Belt) shows a strongly developed foliation

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37 defined by the parallel alignmen t of abundant coarse -grained biotite with less abundant fine-grained sillimanite fibrol ite. Some exposures of the meta-Greyson pelitic schist in the southern study area have a foliation comp rised of biotite + sillimanite fibrolite + muscovite. In the central and eastern study ar ea the pelitic schist grades to a pelitic paragneiss (also correlated with the meta morphosed Greyson Fm.) characterized by a distinct gneissic banding. The gneissic bandi ng in these strata is defined by alternating quartzite and pelitic layers, both typically ranging from ~1 -20 cm in thickness. The quartzite layers of the gneissi c banding are comprised of medium to coarse-grained lightto-dark grey quartzite with minor biotite. Th e pelitic layers are medi um to coarse-grained and comprised largely of (i n increasing relativ e abundance) cordierite + garnet + plagioclase + biotite + K-feldsp ar + quartz + sillimanite fibr olite. Within the pelitic layers very abundant sillimanite fibrolite and much less abundant biotit e (note the lack of muscovite) comprise a strongly developed folia tion parallel to the gneissic banding (Fig. 7-6). The biotite quartzite paragneiss (metamor phosed Ravalli Group) often exhibits a moderately to strongly developed gneissic banding. This gneissic banding is defined by alternating quartzite and more pelitic la yers, both ranging in thickness from a few centimeters to a few meters in some places. Th e quartzite layers are composed of light to medium-grey medium-grained quartzite with mi nor biotite. The more pelitic layers are typically medium-grained and largely compri sed of (in increasing relative abundance) Kfeldspar + biotite (commonly chloritized) + pl agioclase + quartz muscovite. In both the quartzite and more peli tic layers aligned biotite defines a weakly to moderately developed foliation parallel to the gneissic banding (Fig. 7-7a). In addition, in some exposures the

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38 foliation in the more pelitic layers is comprise d of aligned biotite + sillimanite fibrolite muscovite. The calc-silicate paragnei sses correlated with the metamorphosed Helena Formation and Missoula Group (middle and upper Belt, respectively) are also commonly characterized by a prominent gneissic banding. Within these calc-silicate paragneisses, gneissic banding is defined by al ternating light-grey to light-gr een and dark-green layers. The light-grey to light-green layers are mostly composed of quartz + calcite chlorite while the dark-green layers are comprised of diopside + calcite + quartz chlorite tremolite sericite white mica. Light-grey to light-green and dark-green layers are both fine-grained and typicall y range in thickness from 1 mm to ~10 cm (Fig. 7-7b). Aligned chlorite defines a moderately developed folia tion in both of these layer types which is oriented parallel to the gneissic banding. A thin section prepared from the dark green layering in the Missoula Group-e quivalent calc-silicate paragn eiss exhibits highly altered tremolite phenocrysts surrounded by a matrix ri ch in clinopyroxene (Fig. 8). Near the base of both the Helena Formation and Mi ssoula Group the calc-silicate paragneisses grade into layers of alternating biotite schi st and biotite-rich qua rtzite where biotite chlorite define a moderate to strongly de veloped foliation parallel to gneissic banding. The metamorphosed middle Cambrian-equiva lent strata within the LISZ lack gneissic banding but do exhibit variably deve loped metamorphic foliations. The white to pink coarse-grained quartzite correlated w ith metamorphosed Flathead Formation shows a weakly to moderately developed foliation in some areas defined solely by the parallel alignment of muscovite. Fine -grained biotite schist corr elated with metamorphosed

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39 Silver Hill Formation exhibits a moderately developed foliation composed of aligned biotite sillimanite. The white coarse-gra ined marble correlated with the Hasmark Figure 3-6. Outcrop photograph showi ng the strong gneissic banding of the metamorphosed Greyson Formation in th e LISZ. The inset shows a close up of the outcrop exhibiting large garnet porphyroblasts encased by a metamorphic foliation comprised of abunda nt sillimanite fibrolite with less biotite. Photograph is directed to the east.

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40 Figure 3-7. Outcrop photographs of gneissic banding commonly observed in the LISZ. A) Gneissic banding common in the metamorphosed Ravalli Group. Rock hammer handle for scale. B) Gneissic banding common in the metamorphosed Missoula Group. The pe ncil sharpener is for scale. Formation lacks a visible foliation in most places observed. However, a few outcrops exhibit a weak foliation defined by the pa rallel alignment of sparse chlorite. In the central and eastern study area, a pr ominent deformed quartz diorite sill is emplaced within Belt and Middle Cambrian-equivalent metasedimentary strata and

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41 exhibits a moderately to str ongly developed solid-state folia tion. The solid-state foliation is defined by the parallel alignment of el ongate hornblende, K-feldspar, plagioclase, and Figure 3-8. Photomicrograph of the Missoula Group-equivalent calc-silicate paragneiss. Note large altered tremolite grains surrounded by fine-grain ed clinopyroxene in the matrix. Photomicrograph was taken in cross polar light. Trem = Tremolite, cpx = Clinopyroxene. Field of fiew is ~3 mm. quartz (Fig. 3-9). This foliation is concor dant to (parallel with) the metamorphic foliations and gneissic banding in the adjace nt attenuated Belt and middle Cambrianequivalent metase dimentary strata. Mesoscopic-scale folds Mesoscopic-scale folds are found in bot h the meta-Belt and meta-Cambrian metasedimentary strata in the LISZ, but especi ally in meta-Belt metasedimentary strata. These mesoscopic folds typically have ge ometries that range from open (60-120 interlimb angles) to isoclinal (0-10 interlimb angle). The hinge lines of the folds are oriented parallel to the strike of the meta morphic foliation (and/or gneissic banding) in the metasedimentary strata and their plunges vary from steep to shallow. The orientation

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42 of the axial plane with respect to the foliation ranges from up right (70-90) to recumbent (0-10) (e.g., Van Der Pluijm and Marshak, 2004, p. 243-247). In addition, the style of mesoscopic folding in the LISZ ranges from parallel (Class 1B, c onstant layer thickness maintained) to Class 3 type folding (signi ficant thickening in fo ld hinge region and thinning in the limbs, see Ramsay, 1967 for fo ld classification). Figures 3-10 and 3-11 displays common mesoscopic fold geometries found in the Belt metasedimentary strata of the LISZ. It is important to note here that the overall geometry of the mesoscopic folds markedly changes across the east-west striki ng LISZ (and the contact with the deformed quartz diorite sill). Near the center of the LISZ, directly adjacent to the deformed quartz diorite sill, mesoscopic folds are tight to isoclinal and almost exclusively recumbent. This fold geometry is especi ally common in the Greyson pe litic paragneiss adjacent to the sill where the folds so flattened and trans posed it is often difficu lt to distinguish the folds from the planar gneissic banding. Away from the deformed quartz diorite the mesoscopic folds become progressively more open and are typically not recumbent. Mesoscopic-scale boudins Mesoscopic boudins are also common in the Belt-equivalent metasedimentary strata within the LISZ. In the Greyso n Formation and Ravalli Group-equivalent metasedimentary strata these boudins are found in quartzite layers. In the calc-silicate paragneisses correlated with the metamor phosed Helena Formation and Missoula Group the boudins are found within the lighter-colored more quartz-rich layers. Boudins within the Belt-equivalent metasedimentary strata are blocky to tablet or lozenge-shaped and range in length and width from ~5 cm to 1-1.5 m and ~5-25 cm, respectively. In addition, when asymmetrical the longer axes of the boudins are usually oriented parallel

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43 sub-parallel to the metamorphic foli ation (and/or gneissi c banding) in the metasedimentary strata. Figure 3-9. Outcrop photograph of the deform ed quartz diorite sill in the center of the LISZ. The sill exhibits a strong so lid-state foliation (sub-vertical in photograph) concordant to metamo rphic foliation in the adjacent metasedimentary country rocks. This relationship indicates the sill was emplaced into the LISZ prior to or during some phase of solid-state deformation. The photo is directed to the west. Rock hammer handle for scale. As in the case of the mesoscopic folds, the overall geometries of the mesoscopic boudins change in a direction perpendicular to th e overall strike of the LISZ. In the

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44 central LISZ, adjacent to the deformed quartz diorite sill, the boudins are more flattened, lozenge-shaped, and are characte rized by large aspect ratios (i.e., length to width ratio; Fig. 3-12a). However, away from the central LISZ and deformed quartz diorite sill the mesoscopic boudins are typically block-shaped and shorter, having smaller aspect ratios (Fig. 3-12b). Figure 3-10. Outcrop photographs of meso scopic-scale folds found in Belt equivalent metasedimentary strata deformed in the LISZ. A) Nearly isoclinal class 3 folds in meta-Greyson Formation paragneiss. B) Mesoscopic folds in the meta-Ravalli Group biotite quartzite gneiss.

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45 Shear sense The sense of shear for the LISZ is show n by shear-sense indictors found mostly in the meta-Greyson pelitic schist and paragneiss and the deformed quart z diorite sill in the central part of the shear z one. In the meta-Greyson, la rge garnet and cordierite porphyroblasts with pressure shadows tails form these shear-sense indicators. Flattened plagioclase and K-feldspar porphyroc lasts form the shear-sense indicators in the deformed quart diorite sill. In most observations, the porphyroblasts of the metaGreyson and porphyroclasts of the deformed quartz diorite sill formed approximately symmetric sigma-type shear-sense indicators with pressure sh adow tails aligned with the median line of the porphyroblast (the median bisects the porphyroblas t) indicative of pure (or flattening) shear (Fig. 313a, Passchier et al., 1990). Ho wever, in a few observations outcrops located directly east a nd west of the Lake of the Is le show garnet porphyroblasts that form asymmetric sigma-type shear indica tors (with offset pressure shadow tails) showing evidence for left-lateral simple shear (Fig. 3-13b). Description and Distribution of Metamo rphic Phase Assemblages and Textures All Belt and Middle Cambrian-equivalent meta sedimentary strata in the LISZ bear middle to upper-amphibolite facies phase assemblages. The phase assemblages were briefly summarized above in the descriptions of the metamorphic fabrics. Note that a description of these phase assemblages can be found in Appendix B. Here, descriptions and distributions of key or index metamorphic phase assemblages and important textural features are summarized to facilitate a later discussi on concerning the PT history of the LISZ during the Late Cretaceous.

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46 Figure 3-11. Outcrop photographs of meso scopic-scale folds found in Belt-equivalent metasedimentary strata deformed in the LISZ. A) Mesoscopic folds in metaHelena Formation calc-silicate gneiss on the eastern Flank of Mount Tiny. Pencil is for scale. B) Mesoscopic Z-fold found in the meta-Ravalli Group biotite quartzite gneiss. Pencil is for scale. Metamorphic phase assemblages and text ures in the meta-Greyson Formation Pelitic strata correlated with the Greyson Formation exhibit significant changes in metamorphic phases and textures along a transect perpendicular to the strike of the LISZ

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47 and to the contact with the deformed quartz di orite sill. These changes are summarized in the Table 3-1 which includes brief descriptio ns of the metamorphic phase assemblages and important textures for six samples co llected from the metamorphosed Greyson. Figure 3-12. Outcrop photographs of mesosc opic boudins in the LISZ. A) In the outer parts of the shear zone, away from th e contact with deformed quartz diorite sill, mesoscopic boudins are commonly bloc ky. B) In the center of the LISZ, near the contact with the deformed quartz diorite sill boudins are more flat and ellipsoidal.

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48 Samples listed in Table 3-1 are arranged in order of decreasing dist ance from the contact with the deformed quartz diorite sill in the central LISZ. Sample localities are shown in Figure 3-16. In the southernmost part of the study ar ea, away from the contact with quartz diorite sill and the central LI SZ, the pelitic Greyson-equiva lent schist bears abundant coarse-grained biotite, relatively small e uhedral garnets (~1-5 mm in diameter), Kfeldspar, fine-grained sillimanite fibr olite, and primary muscovite (e.g., WG04-108, Table 3-1). This phase assemblage and genera l textural features are maintained within the pelitic Greyson schist north towards the central LISZ until ~1 km from contact with the deformed quartz diorite sill. Here, the pelit ic schist grades to the pelitic paragneiss as described above. At this distance from the si ll, the pelitic paragneiss contains much less abundant biotite, relatively large (~5-10 mm in diameter) inclusion-rich subhedral garnets, K-feldspar, relatively coarser-grain ed sillimanite fibrolite, and no primary muscovite (e.g., WG04-026, Table 3-1). Further north, at a distance of ~0.5 km from the contac t with the deformed quartz diorite sill, some significant changes occur in the phase assemblage and texture of the pelitic paragneiss. Here, the pelitic paragnei ss contains much less biotite, relatively large inclusion-rich subhedral to a nhedral garnets, K-feldspar, ve ry abundant relatively coarsegrained sillimanite fibrolite, large sigmoida l-shaped cordierite porphyroblasts, and no primary muscovite (e.g., ME-231, Table 3-1). In addition, at this distance from the deformed quartz diorite sill, and closer, the pelitic paragneiss is migmatitic and contains granitic leucosome comprised of (in increasing relative abund ant) K-feldspar, quartz, and albite. This leucosome is commonly found in boudin necks within boudinaged quartzite

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49 Figure 3-13. Outcrop photographs of shear sense indicators in the meta-Greyson Formation deformed in the LISZ. A) Symmetric porphyroblast shear sense indicator showing pure (flattening) sh ear. Hand lens for scale, photograph courtesy of Heather Bleick. B) Asy mmetric sigma-type garnet porphyroblast shear sense indicator showing left latera l simple shear sense. Pencil is for scale. Note symmetric shear sense indicators such as shown in A were observed much more frequently than le ft later sense of shear indicator as shown in B. horizons and as thin elonga te pods within the metamorphic foliation and/or gneissic banding of the pelitic paragneiss (Fig. 3-14). In thin section, the leucosome is found within the pressure shadows of garnet and cordierite porphyr oblasts and as thin elongate

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50 pods or lenses within the sillimanite fibrolite -biotite foliation (Fig. 3-15a). In addition, the leucosome is found in thin section as thin fingers or veinlets cross-cutting garnet porphyroblasts extending into the sillimanite fi brolite-biotite foliati on (see the detailed description of ME-231 in cluded in Appendix C). Based on the information summarized above and several other field observations three major mineral/textural zones have b een mapped within the pelitic schist and paragneiss correlated with the metamorphosed Greyson Formation. These three zones are shown on a simplified geologic sketch ma p in Figure 3-16. Zone 1 corresponds to pelitic schist that contains abundant coarse -grained biotite, fine-grained sillimanite fibrolite, relatively small euhedral garnets, and primary muscovite. Zone 2 corresponds to pelitic paragneiss containing, less abundant bi otite, fine-grained sillimanite fibrolite, larger inclusion-rich subhedral to anhedral garnets, and no primary muscovite. Zone 3 represents pelitic paragneiss th at contain little bio tite, relatively abundant coarse-grained sillimanite fibrolite, large inclusion-rich subhe dral to anhedral garn ets, large cordierite porphyroblasts, and no primary muscovite. K-fe ldspar is found in all three mineraltextural zones within the meta-Greyson. Relict and fresh kyanite in the upper Meta-Ravalli Group The uppermost ~100-200 m section of Ra valli Group-equivalent metasedimentary strata mapped is more pelitic than the rest of the section. These more pelitic strata are well exposed near a small unnamed lake at the head of the Twin Lakes Creek drainage in the western LISZ, southwest of Storm Lake just north of the continental divide (Appendix D). Here, the more pelitic upper Ravalli metasedimentary strata contain abundant white to light cream-colored pse udomorphs. In outcrop, these pseudomorphs form elongate blade-shape cross-sections when cut parallel to thei r longer axes. These

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51 elongated pseudomorph cross-section range in length and width from ~1-4 cm and ~2-5 mm, respectively. When cut pe rpendicular to their longer axes, the pseudomorphs form diamond-shaped cross-sections typically ra nging from ~3-10 mm in the longer dimension (Fig. 3-17a). Figure 3-14. Outcrop photographs of me ta-Greyson migmatitic paragneiss showing granitic leucosome commonly found in pressure shadows between quartzite boudins. Upper: Field notebook for scal e; upper photograph is courtesy of Heather Bleick. Lower: Rock hammer is for scale.

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52 Figure 3-15. Photomicrographs of thin sections of the meta-Greyson migmatitic paragneiss exhibiting granitic leuc osome (sample ME-231). A) The leucosome cross-cuts garnet porphyroblast. B) The leucosome is also found in the pressure shadows of large cordierite porphyroblasts. Note the similar placement of the leucosome in pressure shadows at both the outcrop and thin section scale. gt = garnet, crd = cordierite, sill = sillimanite. Photomicrographs were ta ken in cross polar light. In thin section, the pseudomorphs are almo st entirely composed of fine-grained sericite white mica (Fig. 3-17b). Relativel y coarse-grained muscovite laths commonly form corona structures around the pseudomorphs Small relicts of the original mineral

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53 remain within the pseudomorphs as mode rately to highly altered, but high relief fragments that are typically 0.5 mm in diameter or long est dimension (Fig. 3-17b). Based on the blade-shape form (when cut parallel to their longer axes ), the high relief of the internal relict mineral material, and musc ovite corona structures the pseudomorphs exposed at this locality likely re place kyanite. In addition, small ( 1 mm) blade-shaped mineral grains are found scattered thr oughout the matrix betw een large kyanite pseudomorphs (Fig. 3-18). The small bladeshaped grains found in the matrix between the large pseudomorphs are also kyanite, but apparently less altered (Nesse, 1991). Spatial Relationship Between the Lake of the Isle Shear Zone and Late Cretaceous Intrusions The spatial relationship between major Late Cretaceous intrusions and the LISZ is described in some detail in Appendix B. Here a brief summary is given for the important spatial relationships. The LISZ was intrud ed by four major Late Cretaceous intrusions (see map in Appendix F). In the west, the LISZ was intruded by the quartz diorite and granodiorite plutons of the SLS. Both of these intrusions are undeformed and cross-cut the high grade and deformed meta-Belt and me ta-Cambrian metasedimentary strata of the LISZ indicating a post-kinematic relationship to the shear zone. In the east, a quartz diorite pluton (similar to the quartz diorite of the SLS) cross-cuts the Mill Creek nappe indicating a post-kinematic relati onship to the LISZ as well. Between these two areas, a quartz diorite sill was emplaced into the defo rmed metasedimentary strata of the central LISZ. The quartz diorite sill is highly defo rmed and exhibits a well-developed foliation in many places that is concordant to the folia tion in the adjacent metasedimentary country rocks indicating a pre or syn-kinematic rela tionship to the shear zone (see Appendix B and F).

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54Table 3-1. Description and di stribution of key mineral phases and textures in the metamorphosed Greyson Formation. Sample Distance from deformed qd sill (km) Key phases Mineral/textural zone (Fig. 3-16) Textures WG04108 4.0 bt, grt, kfs, ms, sill 1 bt is very abundant and fresh; grt is euhedral, no embayments; no cordierite; fine-grained sill fibrolite WG04095 1.2 bt, grt, kfs, ms, sill 1 possible kyanite relicts; fine-gra ined sill fibrolite; grt is v. euhedral; v. abundant ms WG04101b 1.2 bt, grt, kfs, ms, sill 1 ms present; little sill, but fine -grained variety; bt is v. abundant WG04099 1.2 bt, grt, kfs, ms, sill 1 little sill, fine-grained fibrolite; possible kyanite pseudomorphs; very abundant ms; rock is coarse grained; bt is most abundant; smaller garnets <3 mm in diameter WG04026 0.8 bt, grt, kfs, sill 2 grt is subhedral and inclusion rich; coarser sill fibrolite; lacks ms ME-231 0.4 bt, grt, kfs, sill, crd, leuco 3 bt less abundant and is altered; coarser sill fibrolite is very abundant; larger grt and large crd; crd have abundant fine-grained sill fibrolite inclusions; lacks ms; abundant leuco Note: qd = quartz diorite, bt = biotite, grt = garnet, ms = mu scovite, kfs = K-feldspar, sill = sillimanite, crd = cordierite, and leuco = granitic leucosome.

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55 Figure 3-16. Simplified geologic sketch ma p of the ACC lower plate exposed in the current study area showing the metamorphic mineral/textural zones 1, 2, and 3 mapped in the meta-Greyson Formation adjacent to the deformed quartz diorite sill. The meta-Greyson Formation is brown, the sill is green, and unco rrelated Lower Belt pelitic rocks are light tan. Note the localities of samples discussed in the text and described in Ta ble 3-1 (see text for details).

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Figure 3-17. Kyanite pseudomorphs in the upper meta-Ravalli Group. A) Outcrop photograph of kyanite pseudomorphs. Th is outcrop is located in the western LISZ (see Fig. 3-16 and sample locality for sample WG04-113). Pen for is scale. B) Photomicrograph from a th in section of meta -Ravalli Group sample WG04-113 showing large kyanite pseudomorph made of fine-g rained sericite white mica surrounded by relatively co arse-grained muscovite corona structure. Photomicrograph was taken in cross polar light. Field of view is ~3 mm.

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57 Figure 3-18. Photomicrograph showing fres h kyanite in the upper meta-Ravalli Group. Thin section of WG04-112 shows fine-gra ined fresh kyanite scattered through the matrix of between large kyanite pse udomorphs. Field of view is ~3 mm.

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58 CHAPTER 4 U-PB ZIRCON GEOCHRONOLOGY Purpose and Strategy U-Pb zircon geochronology was employed in this study to provide age constraints on upper-amphibolite facies metamorphism, anatexis, and ductile deformation in the Lake of the Isle shear zone (see Chapter 3). Three samples were chosen from the geochronology. WG04-114 was collected from the undeformed granodiorite of the Storm Lake Stock (SLS) in the western LI SZ (Appendix F). Here, the undeformed granodiorite obliquely crosscuts the deformed metasedimentary stra ta of the western LISZ, indicating that emplacement of the in trusion postdated upper-amphibolite facies metamorphism and ductile deformation in the LISZ. A U-Pb zircon crystallization age from WG04-114 would provide a minimum age limit for upper-amphibolite metamorphism and ductile deformation in the LISZ. Ug-1 was collected from a quartz diorite sill in the central LISZ where the sill is emplaced within a sequence of upperamphibolite grade meta-Belt and meta-Cambrian metasedimentary strata. The quartz diorite sill is deformed and exhibits a well-developed solid-state foliation concordant with the foliation in the adjacent metasedime ntary country rocks. Such a relationship indicates (1) emplacement of the sill pred ated upper-amphibolite metamorphism and ductile deformation in the LISZ or (2) th e sill was emplaced synkinematic to these events. A U-Pb zircon crystallization age fr om the deformed quartz diorite sill sample Ug-1 would provide an upper age limit or direct age constraint for upper-amphibolite facies metamorphism and ductile deforma tion in the LISZ. Finally, sample WG05-02

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59 was also collected from granitic leucosom e found in the migmatitic pelitic paragneiss (metamorphosed Greyson Fm., Lower Belt) in th e central LISZ. Here there is evidence for anatexis during the upper-amphibolite f acies metamorphism and ductile deformation (see Chapter 3). A reliable U-Pb zircon crys tallization age from the leucosome sample WG05-02 would provide a direct age c onstraint on upper-amphibolite facies metamorphism, anatexis, and ductile deformatio n of metasedimentary strata in the LISZ. Relevant Previous U-Pb Zircon Geochronology Prior to this study, there were no U-Pb ge ochronological data av ailable to provide age constraints for the upper-amphibolite f acies metamorphism and ductile deformation in the LISZ. However, Desmarais (1983) re ported U-Pb geochronolog ical data from two deformed granodiorite intrusions in the s outhern Anaconda-Pintlar Range which exhibit solid-state foliations concordant with foliations in metasedimentary country rocks. U-Pb zircon 207Pb / 235U and 206Pb / 238U ages from these deformed intrusions were discordant with upper and lower concordi a intersects at ~ 1780-1890 an d ~ 78-77 Ma, respectively. Desmarais attributed the discordance natu re of the zircons to inheritance from Proterozoic cores and he interpreted the lowe r intersect ages to be minimum ages for emplacement for the intrusions. U-Pb Zircon Geochronology Results Zircons were extracted from samp les WG04-114, Ug-1 and, WG05-02 using standard rock crushing, density, and magnetic se paration techniques. Select zircons were subsequently analyzed along with FC-1 standard zircons (in house 207Pb / 206Pb age = 1086.9 5.3, 207Pb / 235U age = 1091.5 13.4, and 206Pb / 238U age = 1096.7 21.7 Ma) at the University of Florida using lase r ablation multi-collector inductively coupled

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60 plasma mass spectrometry (LA-MC-ICP-MS) Sample preparation, LA-MC-ICP-MS analytical techniques, and data reduc tion are summarized in Appendix A. A common lead correction was applied to th e drift and fractionation corrected U-Pb zircon data obtained from zircons of samples WG04-114, Ug-1, and WG05-02 to evaluate the need for the correction (Appendix A). In the case of WG04-114 and Ug-1, individual U-Pb ages calculated from uncorrect ed and corrected isotopic data were within analytical error indicating negligible concen trations of common lead in these zircons; a common lead correction was deemed unnecessa ry for these analyses. For WG05-02, a common lead correction is necessary due to the antiquity of the zircons and because all of these zircon gave discordant 207Pb / 235U and 206Pb / 238U ages. WG04-114 (Storm Lake Stock Granodiorite) Twenty-eight spot analyses were taken from twenty-two euhedral and inclusion free zircons from sample WG04-114 (average size = 200-300 m), including five paired core and rim analyses from a few larger zi rcons. Individual U-Pb ages for the WG04-114 zircons are summarized in Table 4-1. 207Pb / 206Pb, 207Pb / 235U, and 206Pb / 238U ages are reported for the older zircons. 206Pb / 238U ages are given for younger zircons. Core analyses are indicated by the subscript c in Table 4-1. All ages in Table 4-1 are reported with 2 errors. Three zircon analyses from WG04-114 yielded anomalously old ages (SLS-7c-c, SLS-8a, and SLS-17b-c) These zircons gave discordant 207Pb / 235U and 206Pb / 238U ages ranging from ~1189-239 Ma an d ~599-88 Ma, respectively. However, the 207Pb / 206Pb ages are more consistent and range from ~2502-2243 Ma indicating a Paleoproterozoic or Archean component to these zircons. The remaining twenty-five analyses from WG04-114 gave fair ly consistent Late Cretaceous 206Pb / 238U ages ranging from ~80-71 Ma and are inte rpreted to represent a single magmatic zircon population

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61 related to emplacement of the SL S granodiorite (Table 4-1). Select zircons were (n= 22) are shown plotted on a Tera-Wasserburg diagram and 206Pb / 238U weighted mean zircon age plot in Figure 4-1. Sixteen of these WG04-114 zircons have similar 206Pb / 238U ages and correspond to a 206Pb / 238U weighted mean age of 74.6 0.8 Ma (MSWD = 0.34, 2 ). This age is interpreted to be the age of emplacement for the SLS granodiorite. Ug-1 (The Deformed Quartz Diorite Sill) A total of twenty-four spot analyses we re taken from twenty-four individual euhedral and inclusion free Ug-1 zircons (average size = 200-300 m); core and rim zircon analyses were not distinguished during the analyses. Individua l U-Pb ages for the Ug-1 zircons are summarized in Table 4-2. As the case of WG04-114, 207Pb / 206Pb, 207Pb / 235U, and 206Pb / 238U ages are reported for older zircons and only 206Pb / 238U ages are given for younger zircons. All ages in Table 4-2 are reported with 2 errors. Two Ug-1 zircons yield anomalously old ages (Ug-01_9 and Ug-01_11). These two zircon analyses gave discordant 207Pb / 235U and 206Pb / 238U that range from ~510-471 Ma and ~299-109 Ma, respectively. The remaining twen ty-two Ug-1 zircon analyses yield Late Cretaceous 206Pb / 238U ages of ~77-71 Ma. These zirc ons are interpreted to be of a single magmatic population which crystalli zed during the emplacement of the quartz diorite sill. The twenty-two Late Cretaceous aged Ug-1 analyses are shown plotted on a Tera-Wasserburg diagram and 206Pb / 238U weighted mean zircon age plot in figure 4-2. Sixteen of these zi rcons correspond to a 206Pb / 238U weighted mean age of 75.0 0.8 Ma (MSWD = 0.54, 2 ). This age is interpreted to repr esent the age of emplacement for the quartz diorite sill.

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62Table 4-1. U-Pb LA-MC-ICP-MS analytical results for WG04-114 Radiogenic ratios Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U e.c. SLS-1 0.0573 0.0025 0.0987 0.0066 0.0121 0.0005 0.08 SLS-2 0.0572 0.0009 0.1026 0.0051 0.0121 0.0005 0.11 SLS-3 0.0566 0.0006 0.0891 0.0045 0.0119 0.0005 0.12 SLS-4 0.0546 0.0008 0.0992 0.0062 0.0119 0.0005 0.09 SLS-5a 0.0572 0.0014 0.09300.00470.01160.00050.11 SLS-5b-c 0.0550 0.0017 0.09790.00540.01180.00050.10 SLS-6 0.0739 0.0032 0.12250.00710.01180.00050.07 SLS-7a 0.0614 0.0022 0.09750.00680.01180.00050.08 SLS-7b-c 0.0604 0.0021 0.10510.00720.01140.00050.07 SLS-7c-c 0.1648 0.0011 2. 2235 0.0975 0.0975 0.0045 0.05 SLS-8a 0.1415 0.0064 0.2654 0.0206 0.0137 0.0006 0.03 SLS-8b-c 0.0877 0.0064 0.14340.01410.01170.00050.04 SLS-9-c 0.0554 0.0009 0. 08090.00510.01150.00050.10 SLS-10 0.0517 0.0012 0.09040.00470.01190.00050.11

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63 Table 4-1. Continued. Ages (Ma) Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U SLS-1 77.4 3.4 SLS-2 77.8 3.5 SLS-3 76.5 3.3 SLS-4 76.3 3.4 SLS-5a 74.53.3 SLS-5b-c 75.43.3 SLS-6 75.83.3 SLS-7a 75.43.3 SLS-7b-c 73.13.2 SLS-7c-c 2502 11 1,188.5 30.2 599.9 26.4 SLS-8a 2243 76 239.0 16.4 88.0 3.9 SLS-8b-c 74.93.3 SLS-9-c 73.93.3 SLS-10 76.53.3

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64Table 4-1. Continued. Radiogenic ratios Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U e.c. SLS-11 0.0562 0.00130.09440.00570.01140.00050.09 SLS-12 0.0514 0.00100.10560.00680.01170.00050.08 SLS-13 0.0515 0.00090.09550.00520.01140.00050.10 SLS-14a 0.0517 0.0010 0.0914 0.0058 0.0111 0.0005 0.08 SLS-14b 0.0519 0.0009 0.0973 0.0058 0.0124 0.0005 0.10 SLS-15 0.0535 0.00110.08200.00480.01150.00050.11 SLS-16 0.0798 0.00290.14780.01170.01160.00050.04 SLS-17a 0.0849 0.00630.15480.01640.01170.00050.03 SLS-17b-c 0.1502 0.0031 0.2996 0.0195 0.0140 0.0006 0.03 SLS-18 0.0593 0.00320.09700.00960.01170.00050.05 SLS-19 0.0598 0.00200.09700.00560.01170.00050.09 SLS-20 0.0515 0.0014 0.0880 0.0061 0.0117 0.0005 0.09 SLS-21 0.0540 0.0019 0.0931 0.0072 0.0120 0.0005 0.07 SLS-22 0.1104 0.0034 0.2049 0.0215 0.0123 0.0006 0.03

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65Table 4-1. Continued. Ages (Ma) Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U SLS-11 73.33.3 SLS-12 75.03.4 SLS-13 73.23.3 SLS-14a 71.2 3.1 SLS-14b 79.5 3.5 SLS-15 73.83.3 SLS-16 74.53.3 SLS-17a 74.93.4 SLS-17b-c 2346 37 266.1 15.1 89.4 3.9 SLS-18 75.13.3 SLS-19 75.03.3 SLS-20 75.0 3.4 SLS-21 76.7 3.4 SLS-22 78.8 3.6 Note: e.c. = error correlation between 207Pb/235U and 206Pb/238U errors. Analyses in italics were excluded from the weighted 206Pb/238U mean age calculation.

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66Table 4-2. U-Pb LA-MC-ICP-MS analytical results for Ug-1 Radiogenic ratios Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U e.c. Ug-01_1 0.0429 0.0008 0. 06830.00290.01160.00050.17 Ug-01_2 0.0331 0.0014 0. 05270.00300.01140.00050.17 Ug-01_3 0.0445 0.0008 0. 0679 0.00300.0110 0.0005 0.16 Ug-01_4 0.0506 0.0013 0. 0771 0.00360.0112 0.0005 0.14 Ug-01_5 0.0564 0.0040 0. 0889 0.00760.0113 0.0005 0.06 Ug-01_6 0.0529 0.0022 0. 0824 0.00470.0114 0.0005 0.11 Ug-01_7 0.0482 0.0017 0. 07860.00410.01180.00050.13 Ug-01_8 0.0590 0.0025 0. 09520.00540.01180.00050.09 Ug-01_9 0.0912 0.0006 0. 5902 0.02500.0475 0.0021 0.08 Ug-01_10 0.0616 0.0018 0. 09980.00530.01140.00050.09 Ug-01_11 0.2825 0.0021 0. 6538 0.02840.0170 0.0007 0.03 Ug-01_12 0.0490 0.0004 0. 08020.00330.01160.00050.15 Ug-01_13 0.0678 0.0019 0. 11320.00610.01190.00050.08 Ug-01_13 0.0789 0.0039 0. 13350.00870.01200.00050.06 Ug-01_15 0.0565 0.0020 0. 09520.00540.01180.00050.09 Ug-01_16 0.0560 0.0004 0. 09140.00390.01160.00050.13 Ug-01_17 0.0947 0.0034 0. 16020.00890.01160.00050.06 Ug-01_18 0.0754 0.0041 0. 13360.01030.01160.00050.05 Ug-01_19 0.1107 0.0020 0. 19120.00870.01200.00050.06 Ug-01_20 0.0607 0.0014 0. 10030.00510.01170.00050.10 Ug-01_21 0.0951 0.0031 0. 16100.00860.01190.00050.06 Ug-01_22 0.0561 0.0014 0. 0865 0.00410.0111 0.0005 0.12 Ug-01_23 0.0679 0.0022 0. 1055 0.00560.0113 0.0005 0.09 Ug-01_24 0.0798 0.0038 0. 13290.00900.01170.00050.06

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67Table 4-2. Continued. Ages (Ma) Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U Ug-01_1 74.13.2 Ug-01_2 73.03.1 Ug-01_3 70.7 3.0 Ug-01_4 72.1 3.1 Ug-01_5 72.4 3.1 Ug-01_6 72.9 3.1 Ug-01_7 75.83.3 Ug-01_8 75.53.2 Ug-01_9 1449 13 471.0 15.8 299.1 12.7 Ug-01_10 73.13.1 Ug-01_11 3375 12 510.8 17.3 108.7 4.7 Ug-01_12 74.63.2 Ug-01_13 76.33.3 Ug-01_13 77.03.3 Ug-01_15 75.63.3 Ug-01_16 74.33.2 Ug-01_17 74.43.2 Ug-01_18 74.13.2 Ug-01_19 76.73.3 Ug-01_20 74.83.2 Ug-01_21 76.13.3 Ug-01_22 71.2 3.1 Ug-01_23 72.2 3.1 Ug-01_24 75.23.3 Note: e.c. = error correlation between 207Pb/235U and 206Pb/238U errors. Analyses in italics were excluded from the weighted 206Pb/238U mean age calculation.

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68Table 4-3. U-Pb LA-MC-ICP-MS analytical results for WG05-02 Radiogenic ratios Grain spot 207Pb 207Pb 206Pb 206Pb 235U 238U e.c. L1 0.1027 0.0006 1.20050.06010.10100.00450.08 L2 0.1703 0.0012 1.06160.06780.05880.00290.04 L3 0.1761 0.0010 6.9862 0.2908 0.3003 0.0131 0.05 L4 0.1308 0.0008 1.89210.13170.14650.00680.05 L5 0.1446 0.0009 5.4060 0.2574 0.2714 0.0123 0.05 L6 0.1057 0.0006 3.18370.19490.21970.01260.06 L7 0.1367 0.0008 1.41900.07400.07790.00350.05 L8 0.0988 0.0007 0.37730.03840.04980.00230.06 L9 0.1466 0.0086 0.0346 0.0034 0.0007 0.02 L10 0.1708 0.0010 10.4532 0.4889 0.3830 0.0175 0.04 L11 0.1003 0.0006 1.78370.08480.11180.00500.06 L12 0.2232 0.0019 4.42700.22770.11760.00560.02 L13 0.1074 0.0006 3.33570.15810.19940.00880.06 L14 0.1042 0.0008 1.97340.10560.12150.00610.06 L15 0.1086 0.0008 1.31910.07650.09350.00420.05 L16 0.1870 0.0010 9.9195 0.4345 0.3066 0.0139 0.03 L17 0.1032 0.0006 4.02750.17390.22470.00980.06 L18 0.0983 0.0006 1.10890.06150.07720.00360.06 L19 0.1597 0.0009 3.46820.15180.16440.00730.05 L20 0.1038 0.0006 2.28420.10000.16360.00730.07 L21 0.1634 0.0009 6.3418 0.2750 0.2849 0.0129 0.05 L22 0.1016 0.0006 0.87360.04820.07580.00350.07 L23 0.0950 0.0006 0.46890.02910.04510.00220.08 L24 0.7816 0.0043 14.0487 0.0508 0.0242 0.00 L25 0.1014 0.0006 0.99280.04870.07800.00360.07

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69Table 4-3. Continued. Ages (Ma) Grain spot 207Pb 207Pb 206Pb 206U 235U 238U L1 1671 11800.827.4620.626.5 L2 2559 13734.632.9368.317.7 L3 2616 92109.7 36.3 1693.0 64.8 L4 2105 111078.345.2881.138.2 L5 2280 11 1885.8 40.0 1547.7 62.0 L6 1723 111453.246.21280.366.1 L7 2184 11896.930.6483.920.9 L8 1601 13325.127.9313.414.4 L9 L10 2564 10 2475.8 42.4 2090.5 80.9 L11 1628 101039.530.5683.329.1 L12 3004 131717.441.7716.932.3 L13 1754 111489.536.41172.347.3 L14 1698 131106.535.4739.035.2 L15 1776 13854.133.0576.424.5 L16 2712 92427.3 39.6 1723.8 68.3 L17 1680 101639.834.51306.851.3 L18 1592 12757.629.2479.521.5 L19 2450 91520.033.9981.340.5 L20 1689 101207.430.5976.740.4 L21 2488 92024.3 37.3 1616.2 64.5 L22 1651 11637.525.8471.220.8 L23 1525 20390.419.9284.613.8 L24 L25 1647 11700.124.5484.021.3 Note: e.c. = error correlation between 207Pb/235U and 206Pb/238U errors. Analyses in italics were excluded from the age calculation shown in Figure 4-3b.

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70 Figure 4-1. Tera-Wa sserburg plot and 206Pb / 238U weighted mean zircon age plot for Storm Lake Stock granodiorite sample WG04-114. Only the zircon analyses represented by the darkened ellipses in A and the darkened boxes in B were used in the weighted mean 206Pb / 238U zircon age calculation. U-Pb isotopic data were not corrected for common lead (see text). WG05-02 (Leucosome from th e Meta-Greyson Paragneiss) A total of twenty-five spot analyses were taken from twenty-five individual subhedral and inclusion free zi rcons of sample WG05-02; because of their small size (average size <100 m), exclusive rim analys es were not made from these zircons.

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71 Figure 4-2. Tera-Wa sserburg plot and 206Pb / 238U weighted mean zircon age plot for the deformed quartz diorite sill sample Ug-1. Only the zircon analyses represented by the darkened ellipses in A and the darkened boxes in B were used in the weighted mean 206Pb / 238U zircon age calculation. U-Pb isotopic data were not corrected for common lead (see text). 207Pb/206Pb, 207Pb / 235U, and 206Pb / 238U ages are reported for all twenty-five of the WG05-02 zircon analyses in Table 4-3. Ages reported in Table 4-3 are given with 2 errors. The 207Pb / 206Pb ages range from ~1325-2804 Ma. However, all the WG05-02 zircon analyses gave discordant 207Pb / 235U and 206Pb / 238U ages except for zircon L8. However, zircon L8 is only concordant because of the large errors associated with the

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72207Pb / 235U and 206Pb / 238U ages. The other twenty-f our WG05-02 zircons yielded discordant 207Pb / 235U and 206Pb / 238U ages ranging from ~2476-285 Ma. Figure 4-3a displays a traditional concor dia diagram which includes a ll twenty-five of the WG05-02 zircon analyses. Twelve zircon analyses r oughly defined a discordi a line (see hi ghlighted ellipses). Figure 4-3b displays this discordi a line which intersects concordia at 43 330 Ma and 1736 330 Ma 450 (2 errors). These intersection ages have large errors because no zircon analyses fall near the intersections themselves. As noted, all but one of the WG05-02 leucosome zircons (zircon L8) gave discordant 207Pb / 235U and 206Pb / 238U ages. There are two possible explanations for the largely discordant nature of the leucosome zi rcons: (1) One explanati on is inheritance of older zircon material from the leucosome zircon cores. Because the WG05-02 zircons are small it is entirely possible that both core and rim regions of the zircons were analyzed simultaneously (with beam diam eter of 30-60 m). As a result, the 207Pb / 235U and 206Pb / 238U zircon ages may represent a mixture of two zircon components (i.e., older cores and younger rims). If different propor tions of the two components were analyzed from several zircons then a discordia array or mixing line could form when the isotopic data are plotted on a concordia diagram (e.g., Fi gure 3a). (2) Another explanation for the discordance observed in the WG05-02 leucos ome zircons is lead loss from older, probably Proterozoic or Archean zircons. For example, if very old zircons underwent differential lead loss during a younger isotopic disturbance (e.g., subsequent reheating during metamorphism and/or intrusion) thes e zircons would fall al ong a discordia line on a concordia diagram. In this case, the tw o concordia-discordia in tersects correspond to

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73 the true age (upper inters ect) of the old zircons an d the age of younger isotopic disturbance (lower intersect Faure, 1986; Williams, 1998). Both of these explanations require the presence of Proter ozoic aged or older zircons in the restite from which the WG05-02 leuc osome was partially melted (i.e., the metaGreyson Fm). Ross and Villeneuve (2003) report 207Pb / 206Pb ages from detrital zircons from unmetamorphosed Lower Belt Greyson-equiva lent strata east (in the Helena Salient) and west of the current study area that range from ~1899-1670 Ma; these ages are consistent with the 207Pb / 206Pb ages for the twelve WG05-02 zircon analyses that define the discordia shown in figures 4-3a and b (~1525-1776 Ma, Fig. 4-4). Therefore, it is possible the WG05-02 leucosome incorporated Mesoproterozoic or Paleoproterozoic detrital zircons from the meta-Greys on Formation during high temperature metamorphism and anatexis in the LISZ. Catholuminescence (CL) imaging of the WG05-02 leucosome zircons is needed to dete rmine which of the two above explanations is correct. CL imaging can be used to de termine if the leucosome zircons consist of inherited cores with th in magmatic rims (consistent w ith explanation 1) or lack younger magmatic rims (consistent with explanation 2). The latter implies that no new magmatic rims grew on the older zircon cores during anatexis and formation of the granitic leucosome.

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74 Figure 4-3. Conventional U-Pb concordia plots for zircons for leucosome sample WG0502 collected from the meta-Greyson mi gmatitic paragneiss in the central LISZ. A) Most leucosome zircons gave discordance ages with a 204Pb common lead correction. B) Twelve leucosome zircon analyses fall along a discordia that intersects concordi a at 1736 450 Ma and 43 330. The discordia is most likely th e result of variable lead loss in Paleoproterozoic zircons inherited from the meta-Greyson Formation protolith. Lead loss likely occurred during anatexis and fo rmation of the leucosome itself.

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75 Figure 4-4. 207Pb/206Pb age density plot for the WG05-02 leucosome zircons. 207Pb/206Pb ages range from ~1525-3004 Ma. The leucosome zircons fall into two main age groups: (1) a Mesoproterozoic to Paleoproterozoic group (~1525-1776 Ma) and an (2) Archean group (~2105-3004 Ma). The zircons that form the discordia line shown in fi gures 3-4a and b are part of the Mesoproterozoic to Paleoproterozoic group.

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76 CHAPTER 5 THERMOBAROMETRY Purpose and Strategy As described above, the Lake of the Isle shear zone (LISZ) is a sinuous, middle to upper-amphibolite facies ductile shear zone th at stretches across the exhumed lower plate of the ACC within the curre nt study area (see Appendix F an d description of the LISZ above). Within the LISZ, Mesoproteroz oic Belt Supergroup and middle Cambrianequivalent metasedimentary st rata have undergone major ducti le attenuation (thinning) during upper amphibolite facies metamorphism. As a result, the deformed metasedimentary strata of the LISZ exhibit a strongly transposed metamorphic foliation and common mesoscopic foliation-parallel boudins and near isoclinal folds. In the central LISZ, in the vicinity of the Lake of the Isle, migmatitic pelitic paragneiss (correlated with metamorphos ed Greyson Fm, Lower Belt) bears an uppermost-amphibolite facies phase assemblage and shows evidence for anatexis during the ductile deformation. In this part of th e LISZ granitic leucosome is commonly found in boudin necks and as isolated thin and el ongate pods within the transposed foliation indicating upper-amphibolite f acies metamorphism and anatexis accompanied ductile attenuation of Belt Supergroup and middle Ca mbrian-equivalent metasedimentary strata in the central LISZ. In the eastern part of the study area the LISZ is overpr inted by the Eocene Anaconda mylonite and several listric brittle normal faults relate d to the now largely eroded brittle detachment system (Appendix F). Here, lower to middle greenschist facies

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77 fabrics of the Anaconda gran itoid mylonites contrast sh arply with the plastically deformed upper-amphibolite facies metasedimentary strata of the LISZ which lie structurally beneath the mylonites. This st ructural relationship indicates that the LISZ predated development of the Anaconda mylon ite and brittle detachment system which later facilitated the exhumation of the ACC lower plate. U-Pb geochronological and 40Ar/39Ar thermochronological data obtained from the LISZ in this study give a Late Cretaceous age for the shear zone (ca. ~ 75-74 Ma, see Chapter 4 and 6). These age constraints indicate that upper-a mphibolite facies metamorphism in the LISZ predated tectonic exhumation of the ACC lower pl ate by probably no less than ~23 Ma. The structural relationship between the Anaconda greenschist facies mylonite zone and the LISZ in the eastern study ar ea also shows that the LISZ was exhumed from beneath the brittle detachment fault syst em and ACC upper plate along with the greenschist mylonites during the Eocene. In this study, thermobarometric data were obtained from pelitic strata in the LISZ to constrain the peak pressure-tempera ture associated w ith upper-amphibolite metamorphism, local anatexis, and ductile atte nuation of metasedime ntary strata in the shear zone. Because of the structural rela tionship between the LISZ and the greenschist mylonites pressure constraints from the LI SZ can be used to constrain the maximum amount of exhumation facilitated by extension in the ACC. ME-231 (Migmatitic Meta-Grey son Formation Paragneiss) Sample ME-231 was selected for the therm obarometry needed in this study. ME231 was collected from cordierite-bearing mi gmatitic pelitic paragneiss (meta-Greyson) exposed in the central part of the LISZ, just southwest of the Lake of the Isle (Appendix F). Upper-amphibolite facies metamorphism was most intense here and the meta-

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78 Greyson strata bear a phase assemblage of (i n increasing relative a bundance) cordierite + garnet + K-feldspar + albite plagioclase + bi otite + quartz + sillimanite fibrolite. This metamorphic assemblage is indicative of uppermost amphibolite facies in the second sillimanite-K-feldspar zone (Spear, 1993). A de tailed description of ME-231 is included in Appendix C. Relevant Previous Pressure-temperature Constraints Prior to this study, quantitative thermobarometric data were not available from the metasedimentary strata deformed in the LISZ or any other part of the Anaconda-Pintlar and Flint Creek Ranges. A few previous st udies do document amphibolite facies regional metamorphism of metasedimentary rocks in th e ACC lower plate. The observations and pressure-temperature estimates made in these studies are important to later discussion of regional metamorphism in the ACC lower plate. Desmarais (1983) described amphibolite faci es pelitic and calcsilicate schists and paragneisses in ACC lower plate exposed in the southern Anaconda-Pintlar Range, southwest of the current study area. He descri bed pelitic strata with an upper-amphibolite facies assemblage of quartz + biotite + plagioclase + K-feldspar + sillimanite muscovite. Desmarais correlated these strata with the Mount Shields Formation (middle Missoula Group equivalent, upper Belt). He also observed migmatitic paragneisses in some areas suggesting that local anatex is accompanied upperamphibolite facies metamorphism. Based on these observations, Desmarais estimated peak metamorphic conditions in the southern Anaconda-Pintlar range reached pressures of ~3.0-7.0 kbar and temperatures of ~600-700C. Flood (1974) mapped and described the Fish trap Creek nappe, a large west-verging recumbent nappe fold structure exposed in the southern Anaconda-Pintlar Range. The

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79 Fishtrap Creek nappe is comprised of metase dimentary strata he correlated with (from structurally lowest to highest) the Prichard Formation (Greyson Fm. equivalent), Ravalli Group, and the Wallace Formation (Helena Fm. e quivalent). Flood described these metaBelt units as biotite-muscovite schist, quartzite, and calc-si licate schist, respectively. Flood (1974) also documented and measured well-developed metamorphic foliation throughout the Fishtrap Creek nappe and commo n mesoscopic-scale folds. Based on the metamorphic assemblages of these Beltcorrelated metasedimentary strata Flood estimated that peak metamorphic conditions in the southernmost Anaconda-Pintlar Range reached pressures of ~2-4 kbar and temperatures of ~550-650C. Stuart (1966) mapped and described metase dimentary strata adjacent to the Late Cretaceous Royal Stock granodiorite in the no rtheastern Flint Creek Range, north of the current study area. He described a some what narrow contact aureole surrounding the Royal Stock superimposed on more widespr ead and regional metamorphic fabrics. Within the Royal Stock contact aureole he notes randomly oriented andalusite porphyroblasts overprinting the regional me tamorphic fabric. Outside the contact aureole, Stuart documented highly deform ed metasedimentary strata comprised of muscovite + biotite + quartz + co rdierite rather than andalusite In thin section, he noted large sigmoidal shaped cordierite porphyroblasts with their long axes oriented parallel to the pervasive regional meta morphic foliation. These obser vations indicate an earlier regional metamorphic event at middle to uppe r-amphibolite facies followed by a lower pressure event associated with the intrusi on of the Royal Stock. Notably, Stuart also documented a general increase in the regional metamorphic gr ade west to east across the northern Flint Creek Range, ranging from greenschist to upper-amphibolite facies.

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80 Thermobarometry Results Pressure-temperature estimates were obtaine d from polished thin sections prepared from sample ME-231 by electron microprobe analyses and subsequent thermodynamic calculations made using the computer pr ograms AX (activity-c omposition) by Holland and Powell (2000) and THERMOCALC v. 3.21 by Powell et al. (1998). For comparison, a pressure-temperature estimate for ME-231 wa s also made using the computer program Geothermobarometry (GTB) v. 2.1 by Spear and Kohn (1999). Sample preparation, microprobe instrumentation and analytical pr ocedures are summarized in Appendix A. Thermodynamic calculations made using AX, THERMOCALC, and GTB are discussed below. Electron Microprobe Analyses Major elemental compositions were obtai ned from individual mineral phases in sample ME-231 using a JOEL Superprobe elect ron microprobe at Florida International University. The results from the analyses of garnet, biotite, al bite plagioclase and cordierite are summarized in Table 5-1. Each mineral analysis is reported in elemental oxide weight percent and element oxide totals from each analysis are shown in the far right column of the table. Garnet and cord ierite analyses labeled with the suffixes c and r, indicate mineral core and rim analyses, respectivel y. Garnet and cordierite analyses without these suffixes were taken from intermediate spots on the mineral, between the rim and core region. All biotite and plagioclase analyses from sample ME231 are rim analyses. Garnet A total of thirty-five spot analyzes were taken from large and small garnet porphoroblasts. Most of these analyses gave elemental oxides totals of 100 1 weight

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81Table 5-1. Results from electr on microprobe mineral analyses. Analyses SiO2 TiO2 Cr2O3 Na2O K2O Al2O3 MnO FeO MgO CaO Total Garnet PMT-2-1-gt1c 36.94 0.23 0.170.32 0.1021.291.06 36.932.541.02100.60 PMT-2-1-gt2 37.36 0.15 0.170.37 0.1021.461.29 36.182.971.03101.06 PMT-2-1-gt3 37.02 0.12 0.140.33 0.1021.331.28 36.302.991.09100.70 PMT-2-1-gt4 37.00 0.23 0.250.31 0.1121.501.04 36.622.990.94100.97 PMT-2-1-gt5 36.72 0.21 0.150.28 0.1121.290.58 37.852.241.14100.57 PMT-2-3-gt1 36.68 0.05 0.060.01 0.0121.050.33 37.722.471.2299.61 PMT-2-3-gt2 36.42 0.07 0.070.00 0.0121.290.44 37.502.560.9099.26 PMT-2-3-gt3 36.51 0.00 0.060.02 0.0021.190.40 36.592.501.1298.40 PMT-2-3-gt4 36.41 0.00 0.050.00 0.0121.360.36 36.862.550.9598.54 PMT-2-6-gt1 36.37 0.00 0.040.00 0.0021.080.29 38.312.130.8299.03 PMT-2-6-gt2 36.12 0.00 0.020.00 0.0021.130.49 34.292.251.1995.50 PMT-2-6-gt3 36.79 0.00 0.000.00 0.0121.070.49 37.682.181.0099.21 PMT-3-1-gt1 36.81 0.13 0.150.30 0.1321.570.51 38.352.911.10101.94 PMT-3-1-gt2 36.73 0.13 0.160.31 0.1221.391.13 36.783.110.95100.81 PMT-3-1-gt3 36.60 0.18 0.140.29 0.1321.241.40 36.752.891.02100.64 PMT-3-1-gt4 36.17 0.13 0.160.29 0.1221.140.95 36.642.941.0499.57 PMT-3-1-gt5 36.58 0.17 0.120.29 0.1121.380.49 37.962.931.26101.28 PMT-3-1-gt6 36.70 0.21 0.330.30 0.1121.130.76 37.252.931.45101.16 PMT-3-1-gt7 36.18 0.12 0.340.27 0.1221.170.81 36.842.541.3099.70 PMT-3-1-gt8 36.37 0.08 0.160.29 0.1320.930.63 37.502.971.25100.30 PMT-3-1-gt9 36.54 0.19 0.260.31 0.1221.580.99 36.942.961.29101.17 PMT-3-1-gt10 36.38 0.21 0.210. 300.1221.130.99 37.093.011.23100.67

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82Table 5-1. Continued. Analyses SiO2 TiO2 Cr2O3 Na2O K2O Al2O3 MnO FeO MgO CaO Total PMT-3-2-gt1-1 37.87 0.00 0.020. 000.0021.410.37 36.612.620.9799.86 PMT-3-2-gt1-2 37.51 0.01 0.010. 000.0221.340.31 37.332.541.12100.18 PMT-3-2-gt1-3 37.52 0.02 0.050. 020.0021.210.32 38.362.111.00100.60 PMT-3-2-gt2-1 37.82 0.00 0.000. 030.0020.930.38 37.032.500.9399.62 PMT-3-2-gt2-2 37.79 0.06 0.000. 000.0120.880.33 36.712.371.2399.38 PMT-3-2-gt2-3 37.84 0.00 0.020. 020.0021.070.36 36.832.311.0099.45 PMT-3-3-gt1 37.23 0.19 0.180.27 0.1021.190.49 37.642.901.19101.38 PMT-3-3-gt2 36.47 0.09 0.150.31 0.1121.400.50 36.992.900.9799.89 PMT-3-3-gt3 36.60 0.18 0.130.32 0.1021.460.54 37.412.871.16100.76 PMT-3-3-gt4 36.38 0.16 0.190.30 0.1121.520.54 37.442.901.15100.68 PMT-3-3-gt5 36.84 0.12 0.180.29 0.1021.330.46 38.062.971.05101.39 PMT-3-3-gt6 36.91 0.12 0.140.31 0.1121.160.53 37.572.890.96100.69 PMT-3-3-gt7 36.95 0.15 0.210.32 0.1021.320.54 37.302.901.24101.01 Biotite PMT-2-3-bt1 33.48 2.08 0.120.16 8.6620.610.00 24.944.520.0094.56 PMT-2-3-bt2 32.09 3.48 0.110.21 8.5418.940.03 23.473.770.1190.75 PMT-2-3-bt3 32.92 3.34 0.140.22 8.5119.860.08 24.634.410.0194.11 PMT-2-6-bt1 33.71 4.25 0.180.27 8.7119.250.00 22.795.160.0094.31 PMT-2-6-bt2 33.21 3.74 0.250.13 8.5219.320.01 23.395.560.0094.12 PMT-2-6-bt3 33.01 3.79 0.100.21 8.7419.500.00 22.485.190.0093.02 PMT-2-6-bt4 33.01 3.36 0.110.23 8.4618.960.00 22.464.990.0191.59 PMT-2-6-bt5 32.61 2.55 0.170.23 8.0719.420.07 23.025.230.0191.38

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83Table 5-1. Continued. Analyses SiO2 TiO2 Cr2O3 Na2O K2O Al2O3 MnO FeO MgO CaO Total PMT-2-6-bt6 32.58 3.92 0.060.25 8.4718.960.05 21.255.060.0190.61 PMT-2-6-bt7 32.74 4.46 0.100.23 7.9319.290.00 22.464.820.0092.04 PMT-3-1-bt1 33.50 3.40 0.230.43 8.3819.370.12 23.845.110.1594.53 PMT-3-1-bt2 32.89 3.80 0.300.48 8.4319.070.14 24.435.260.1594.95 PMT-3-1-bt3 32.93 3.55 0.200.42 8.1319.280.13 23.434.980.1793.22 PMT-3-1-bt4 33.87 2.82 0.250.43 8.2319.520.16 24.115.390.1694.95 PMT-3-1-bt5 33.64 2.62 0.260.42 8.0520.060.19 24.155.230.1594.77 PMT-3-1-bt6 34.13 3.27 0.240.41 8.3020.030.13 23.644.810.1495.09 PMT-3-1-bt7 33.37 3.31 0.220.47 8.2319.850.21 23.905.070.1694.79 PMT-3-1-bt8 33.63 3.58 0.280.48 8.1019.260.20 23.235.540.1794.45 PMT-3-1-bt9 34.08 3.42 0.280.42 8.1919.500.16 23.625.850.1495.66 PMT-3-1-bt10 33.30 3.67 0.280. 407.9919.060.14 24.175.960.1595.10 PMT-3-2-bt1 33.92 2.02 0.060.21 8.7020.850.07 24.194.580.0194.60 PMT-3-2-bt2 34.05 2.85 0.150.20 8.6020.050.00 24.034.550.0094.47 PMT-3-2-bt3 33.78 2.58 0.050.18 8.5419.830.07 24.484.840.0194.35 PMT-3-2-bt4 33.67 2.84 0.200.16 8.9319.460.00 23.564.970.0093.81 PMT-3-2-bt5 34.08 2.52 0.130.16 8.6119.720.00 23.625.070.0093.91 PMT-3-2-bt6 33.58 3.14 0.100.19 8.8119.980.02 22.825.060.0193.71 PMT-3-3-bt1 33.54 3.22 0.290.44 7.9719.490.15 24.405.500.1495.14 PMT-3-3-bt2 33.55 3.50 0.250.41 8.2619.320.12 23.595.290.1394.42 PMT-3-3-bt3 33.42 3.55 0.340.50 7.9619.000.11 24.265.300.1694.60 PMT-3-3-bt4 34.43 3.03 0.240.52 8.1819.970.13 24.655.270.1596.58 PMT-3-3-bt5 31.31 2.54 0.190.37 6.6019.800.13 25.025.590.1791.73

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84Table 5-1. Continued. Analyses SiO2 TiO2 Cr2O3 Na2O K2O Al2O3 MnO FeO MgO CaO Total Plagioclase PMT-3-3-bt6 33.74 3.06 0.240.39 7.7920.380.11 24.595.050.1595.50 PMT-3-3-bt7 34.46 3.47 0.310.38 8.0619.720.11 23.095.720.1495.46 PMT-3-3-bt8 34.11 2.93 0.360.47 7.8720.170.15 24.125.950.1596.28 PMT-2-6-pl1 59.96 0.00 0.016.28 0.0624.670.00 0.050.006.0797.10 PMT-2-6-pl2 60.94 0.02 0.005.19 0.0824.610.00 0.030.005.9196.77 PMT-2-6-pl3 57.90 0.06 0.007.34 0.0624.800.02 0.010.007.0497.22 PMT-2-6-pl4 60.05 0.00 0.007.04 0.0824.410.00 0.000.003.8595.44 PMT-2-6-pl5 59.46 0.03 0.006.81 0.0624.740.03 0.000.006.2597.38 PMT-3-1-pl1 58.74 0.14 0.187.18 0.1324.760.10 0.280.267.5099.26 PMT-3-1-pl2 58.57 0.08 0.077.18 0.1724.650.11 0.420.277.5899.10 PMT-3-1-pl3 59.38 0.19 0.217.39 0.1424.170.09 0.320.267.0599.18 PMT-3-1-pl4 58.16 0.10 0.187.00 0.2024.700.07 0.490.287.8699.04 PMT-3-1-pl5 59.02 0.05 0.137.12 0.1324.540.10 0.340.287.3899.07 PMT-3-1-pl6 59.83 0.11 0.117.43 0.1524.230.13 0.310.297.0299.61 PMT-3-1-pl7 59.36 0.09 0.097.17 0.1524.790.08 0.250.277.2899.51 PMT-3-1-pl8 60.35 0.11 0.147.63 0.1323.990.09 0.310.286.5399.55 PMT-3-2-pl1 61.02 0.00 0.057.38 0.0623.970.01 0.230.006.4199.12 PMT-3-2-pl2 60.55 0.00 0.007.52 0.0524.110.00 0.130.006.4798.84 PMT-3-2-pl3 60.94 0.00 0.087.70 0.0623.950.00 0.220.006.4099.35 PMT-3-2-pl4 61.86 0.00 0.015.39 0.0524.340.02 0.180.006.4998.35 PMT-3-3-pl1 59.52 0.00 0.106.37 0.0425.470.05 0.110.008.1099.75 PMT-3-3-pl2 58.89 0.00 0.006.12 0.0325.180.03 0.090.008.3098.64

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85Table 5-1. Continued. Analyses SiO2 TiO2 Cr2O3 Na2O K2O Al2O3 MnO FeO MgO CaO Total PMT-3-3-pl2 60.77 0.03 0.057.66 0.0524.150.00 0.310.006.5299.52 PMT-3-3-pl4 58.82 0.00 0.016.80 0.0325.130.04 0.290.028.1299.26 PMT-3-3-pl5 60.29 0.00 0.007.35 0.0424.210.01 0.270.007.1099.26 PMT-3-3-pl6 59.55 0.00 0.056.87 0.0224.780.00 0.110.007.7899.15 PMT-3-3-pl6 59.42 0.00 0.006.75 0.0324.970.03 0.120.007.8699.18 PMT-3-3-pl7 60.23 0.00 0.007.25 0.0523.690.02 0.050.006.8898.17 PMT-3-3-pl8 58.87 0.00 0.006.86 0.0524.350.01 0.290.077.0797.58 PMT-3-3-pl9 58.69 0.04 0.005.76 0.0524.150.03 0.440.006.8195.97 Cordierite PMT-2-3-crd1 47.78 0.04 0.000. 170.0031.020.07 12.165.230.0396.50 PMT-2-3-crd2 48.54 0.04 0.000. 140.0131.350.07 11.136.290.0397.60 PMT-2-3-crd3 45.90 0.00 0.000. 160.0132.550.00 11.456.300.0296.39 PMT-2-8-crd1 47.09 0.09 0.210. 400.0931.950.09 11.486.530.1598.08 PMT-2-8-crd2 46.90 0.10 0.100. 380.1032.280.15 10.766.840.1397.74 PMT-2-8-crd3 47.02 0.09 0.120. 350.0931.730.09 11.666.460.1397.74 PMT-3-4-crd1 49.37 0.13 0.150. 370.0831.750.17 11.056.640.1399.85 PMT-3-4-crd2 48.78 0.09 0.090. 360.0831.320.13 11.026.910.1598.92 PMT-3-4-crd3 48.98 0.11 0.150. 340.0931.210.16 11.226.700.1599.10 Note: gt = garnet, bt = biotite, pl = plagioclase, crd = cordie rite. The letter c indicates a core analyses. All analyses are reported in weight percent elemental oxides.

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percent, indicating high quality microprobe analyses. Ten garnet analyses did fall slightly outside this range, total to 100 1.5 weight percent. A Fe composition map were made from one large garnet in ME-231 s how that ME-231 garnets are Fe-rich and homogeneous with respect to Fe (i.e., no ev idence for composition zoning, Appendix F). Biotite Thirty-four spot analyses were taken fr om fine-grained biotite of the ME-231 matrix. On average, these microprobe analys es totaled to 94.1 elemental oxides weight percent. The lower total elemental oxide weight percents of these analyses are probably due to significant quantities of H2O in the biotites of ME231. For these analyses the JOEL microprobe was not calibrated to measure H2O and therefore did not measure H2O. However, most biotites c ontain 5-6 weight percent H2O within their crystalline structures (Deer et al. 1992). Therefore, biot ites with element oxide weight percent totals 94.0 are considered reliable here. Plagioclase Fine-grained plagioclase was analyzed fr om the matrix of ME-231. A total of twenty-seven spot analyses were taken from areas in the polished thin sections with abundant garnet, quartz and sillimanite. Sixt een of these analyses fell within 100 1 total elemental oxide weight percent are considered reliable. The other eleven plagioclase analyses gave total elementa l oxide weight percen ts < 99.0 and are not considered reliable and were not used in subsequent calculations. Cordierite Nine spot analyses were taken from th ree large cordierite porphoroblasts in ME231. All cordierite analyses totaled to < 100 weight percent elemental oxides. However, two analyses (3-4-crd1 and 3-4-crd3) from one cordierite gave elemental oxides total

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87 weight percents of 100 1, indicating th ese were good analyses. The other seven cordierite analyses yielded elemental oxi des total weight percent of 96.4-98.9. Lower elemental oxides total weight percents in these analyses may be due to undetected H2O in the cordierite, as in the case of the biotites (Nesse, 1991). Only cordierite analyses 3-4cord1 and 3-4-cord3 were used in subs equent thermodynamic calculations. Mineral End Member Activity Calculations made using AX A total of six garnet, twenty-two biotite, eighteen albite plagioclase, and two cordierite robust rim microprobe analyses from one of the polished thin sections were individually averaged and then imported into the computer pr ogram AX (Table 5-2). The AX program was used to calculate mineral end member activities and to convert the elemental oxide data to cation unit formulas needed for subsequent thermobarometric calculations made using THERMOCALC. End member activities and cation unit formulas were calculated in AX at 4.0 kbar and 800C, an approximate PT estimate for sample ME-231 based on its metamorphic petrology. End member activities for each mineral phase were calculated using id eal and non-ideal so lution mixing models incorporated within AX. Cation unit formul as were calculated with ferric iron (Fe3+) estimation. The AX calculated end member activity and cation unit formulas are summarized in the AX output file included in Appendix E. See Holland and Powell (1998) or http://www.earthsci.unimelb.edu.au/tpg/thermocalc/ for a more detailed description of the co mputer program AX. Pressure-temperature Estimates using THERMOCALC Mineral end member activities calculated in AX were imported into the computer program THERMOCALC to calculate aver age PT estimates for the ME-231 phase assemblage. THERMOCALC utilizes the intern ally consistent thermodynamic database

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88 Table 5-2. Averaged ME-231 electron micropr obe analyses used in thermobarometry Mineral grt rims bt rims pl rims crd rims n 6 22 18 2 SiO2 36.88 33.76 59.62 49.17 TiO2 0.14 3.12 0.05 0.12 Al2O3 21.32 19.71 24.54 31.48 Cr2O3 0.16 0.23 0.08 0.15 Fe2O3 2.71 0.00 0.28 1.33 FeO 35.37 23.93 0.00 9.94 MnO 0.50 0.11 0.05 0.17 MgO 2.92 5.24 0.12 6.67 CaO 1.13 0.11 7.29 0.14 Na2O 0.30 0.37 7.15 0.35 K2O 0.11 8.28 0.09 0.09 Totals 101.56 94.88 99.27 99.61 Oxygens 12.00 11.00 8.00 18.00 Si 2.93 2.63 2.68 5.06 Ti 0.01 0.18 0.00 0.01 Al 2.00 1.81 1.30 3.82 Cr 0.01 0.01 0.00 0.01 Fe3+ 0.16 0.00 0.01 0.10 Fe2+ 2.35 1.56 0.00 0.86 Mn 0.03 0.01 0.00 0.01 Mg 0.35 0.61 0.01 1.02 Ca 0.10 0.01 0.35 0.02 Na 0.05 0.06 0.62 0.07 K 0.01 0.83 0.01 0.01 Cation sum 8.0 7.7 5.0 11.0 XMg 0.1 XFe 2.6 XNa 1.8 XMg 1.1 XCa 0.0 XMg 0.4 XCa 0.6 XFe 0.9 XFe 5.3 XMn 0.0 Note: grt = garnet, bt = biotite, pl = plagioclase, crd = cordierite, n = number of individual analyses averaged. HP98 to make pressure-temperature estimates for rock phase assemblages based on imported mineral end member activities (e .g., Holland and Powell, 1998). For the ME-

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89 231 phase assemblage PT estimates made he re, the activities of sillimanite and quartz were assumed to be 1.0 by convention because these phases are considered pure mineral end members. Using THERMOCALC, a set of five independent reactions was calculated from the imported ME-231 mine ral end member activities: 1 grossular + 1 quartz + 2 sillimanite 3 anorthite (5-1) 2 pyrope + 5 quartz + 4 sillimanite 3 Mg-cordierite (5-2) 2 almandine + 5 quartz + 4 sillimanite 3 Fe-cordierite (5-3) 1 pyrope + 1 eastonite + 3 quartz 1 phlogopite + 1 Mg-cordierite (5-4) 1 almandine + 1 eastonite + 3 quartz 1 annite + 1 Mg-cordierite (5-5) Reaction 5-1 is the GASP, geobarometer, which utilizes the equilibrium between garnet, aluminosilicate, quartz, and plagiocl ase to estimate pressure s from pelitic rock phase assemblages (e.g., Koziol and Newton, 1988; Spear, 1993). The GASP geobarometer is a net transfer reaction, wh ich means phases are consumed and produced across the reaction (Spear, 1993). The GASP act s as an adequate geobarometer because it is sensitive to volume change (large V) and less sensitive to temperature (i.e., small S and H). Reactions 5-2, 5-3 and 5-4 are also net transfer reactions that act as geobarometers for the ME-231 phase assemblage. Reaction 5 is an exchange reaction, meaning chemical components are exchange d between the phases across the reaction without net consumption or production of any phases (Spear, 1993). In the case of reaction 5-5, iron and magnesium are exchanged between garnet, biotit e and cordierite as a function of temperature. The five independent reacti ons calculated in THERMOCALC have uncertainties in PT space, where these uncertainties are a f unction of the uncertainties in electron

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90 microprobe analyses, calculated mineral end member activities, and modeled thermodynamic data, especially enthalpy ( H, Holland and Powell, 1998). THERMOCALC uses the least-squares method (e.g., York, 1969) to vary the positions of the independent reactions in PT space relative to these uncertainties and to reaction correlations (reactions with common end members are correlated in PT space) until a statistical intersection occurs at one point in PT space, the so-called average PT (e.g., Powell and Holland, 1994). In the case of ME-231, the five independent reactions intersect at an average PT of 3.8 1.8 kbar and 657 176C (errors are 2 ; also see the THERMOCALC output file in Appendix E). The average PT estimate is plotted in Figure 5-1a on a phase diagram modified from Spear et al. (1999, see their Fig. 2, p. 19) representing the KFMASH (K2O-FeOMgO-Al2O3-SiO2-H2O) chemical system for pelitic rocks. The average PT estimate is marked by the star on the phase diagram in Figure 5-1a. In add ition, the pressure and temperature errors associated with the av erage PT estimate are represented by the two sigma error ellipse shown in figure X; the pos ition and shape of th is error ellipse is defined by the statistical intersection of the fi ve independent reactions calculated for the ME-231 phase assemblage in PT space. The average PT estimate made here provide s moderate constraints on the pressure and relatively poor constraints on temperature for the phase assemblage of ME-231 (Fig. 5-1a). Only the high temperature portion of the average PT error ellipse actually falls within the divariant field consistent with the metamorphic petrology of ME-231 (see description of ME-231 in Appendix C). The la rge uncertainties in th e temperature of the average PT estimate can be explained by the re lative positions of the five independent

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91 Figure 5-1. Phase diagrams showing PT estimates for ME-231, migmatitic Greyson paragneiss from the Lake of the Isle shear zone. A) THERMOCALC average and locus PT estimates. Uncertainti es in the THERMOCALC average and locus PT estimates are represent by the dashed ellipse and box, respectively. The intersection (darkened region) of these two PT calculates is the best PT estimate for ME-231. B) Geothermobarometry (GTB) PT estimate for ME231. The best GTB PT estimate is de fined by intersections of several independently determined garnet-b iotite exchange and GASP reaction calibrations. In both A and B Al2SiO5 stability fields are from Spear (1993) and univariant reaction curves are taken from Spear et al. (1999).

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92 reactions in PT space. Figure 5-2 displays a pressure-temperature plot showing the five independent reactions calculated from the mi neral end member data for ME-231 plotted in PT space. Considerable spread (separat ion) between the five independent reactions does not permit a well-defined statistical inters ection to be made between the reactions. Consequently, a precise average PT estimat e cannot be made for the ME-231 using the intersection of these five independent reactions. Figur e 5-2 also explains why temperature is less constrained than pressu re for the ME-231 phase assemblage. Because all five independent reactions are characterized by gentle to moderate Clapeyron slopes (the reaction curve slope in PT space), the temperature is much less constrained than pressure. A more accurate PT estimate can be made for the ME-231 phase assemblage using the average locus option in THERMOCALC rather than the average PT option (e.g., Powell and Holland, 1994). The average locu s option allows the user to calculate pressure or temperature independently from the one another, provided a geologically reasonable range of values can be estimated for one of the unknowns (see Powell and Holland, 1994 for a description of the aver age PT and locus pressure-temperature estimates). Since the range of possible temperatures consistent with the ME-231 assemblage is fairly well constrained to ~ 750-900C (Fig. 5-1a), th e average pressure can be calculated in THERMOCALC using the average locus op tion. Pressures calculated over these temperatures range from 4.6-6.0 kbar; however locus pressures above ~830900C fall outside of the divariant field of the ME-231 phase assemblage and are therefore not considered realistic (i.e., these pressures fall within the granulite facies). Notably, locus pressures from ~750-825C overlap with the high temperature side of the

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93 Figure 5-2. Phase diagram showing phase equilibria used by THERMOCALC to calculate the average PT estimate fo r sample ME-231. The poorly defined intersection between these reactions in PT space resulted in an average PT estimate with large uncertainties, especially in temperature. average PT error ellipse within the central part of the ME-231 divariant field (Fig. 5-1a). This region of overlap (darkened region) is considered the best (i.e., the most statistically sound) PT estimate for sample ME-231, corresp onding to a pressure and temperature of 3.2-5.3 kbar and 750-825C, respectively. Pressure-temperature Estimates us ing Geothermobarometry (GTB) To obtain a pressure-temperature estim ate for ME-231 using GTB, the same averaged weight percent elemental oxide data (see Table 5-2) were first converted to cation unit formulas using the Microsoft Excel based spreadsheet pr ogram Formula.xls; this program is available at http://www.earth.ox.ac.uk/~d avewa/pt/th_tools.html The calculated cations for garnet, biotite, and plagioclase were then manually imported into GTB. GTB uses the imported cation data to calculate mineral end member activities then

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94 the positions of several geothermometer a nd geobarometer reactions in PT space; the intersection of these reactions de fines a pressure-temperature es timate. It is important to note GTB uses phase equilibria calibrations from several different sources to make a pressure-temperature estimate (e.g., the geot hermometer and geobarometer equilibria are calibrated by different worker s, and different thermodynamic data). Therefore, the GTB pressure-temperature estimates are not inter nally consistent as pressure-temperature estimates made using the THERMOCALC. Figure 5-1b shows the positions of the Fe-Mg garnet-biotite exchange geothermometer: 1 almandine + 1 phlogopite 1 pyrope + 1 annite (5-6) and the garnet-aluminosilicate-quartz-plagioclase (GASP) geobarometer (see reaction 5-1 above) determined using several different calibrations. The re gion of intersection between the garnet-biotite exchange geothe rmometer and GASP geobarometer (darkened region), as defined by the diffe rent calibrations, corresponds to a PT estimate for ME-231 of ~3-6 kbar and ~725-850C (Fig. 5-1b). Thes e GTB pressure-temperature estimate for ME-231 is in good agreement with the THER MOCALC best PT estimate. The GTB pressure-temperature estimate also corre sponds to upper-amphibolite conditions uptemperature of the minimal pelite melt curve (Fig. 5-1b).

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95 CHAPTER 6 40AR/39AR THERMOCHRONOLOGY Purpose and Strategy Mineral cooling ages obtained by 40Ar/39Ar thermochronology can be used to constrain the cooling and exhumation histor ies of highly extended terranes such as metamorphic core complexes (e.g., Foster a nd John, 1999; Stockli, 2005). Rock samples are typically collected along transe cts in the lower plate parallel to the direction of slip on the bounding detachment fault zone(s). Appare nt cooling ages are then obtained from individual minerals separates extracte d from the transect samples using the 40Ar/39Ar method. These data can then be used to esti mate the timing of the onset and duration of extension, rate(s) of ex tension (e.g., slip rates on detachment faults), cooling histories for lower plates, detachment geometry, and amount s of vertical and hor izontal displacement along detachment fault systems (e.g., Foster et al., 1993; John and Foster, 1993; Scott et al., 1998; Foster and John, 1999; Br ichau et al., 2005). In order to constrain such aspects of extension in the ACC a suite of sixteen rock samples was collected for 40Ar/39Ar thermochronology along a transect across the lo wer plate, parallel to the approximate direction of slip (ESE, ~105) on the east-bound ing detachment fault system. The 40Ar/39Ar thermochronology transect of th is study was carried out in the northeastern Anaconda-Pintlar Ra nge, southwest of the town of Anaconda, MT. Here, upper plate rocks have been almost entirely removed, exposing the metamorphic-plutonic lower plate of the ACC for a lateral (eastwest) distance greater than 20 km. Rock samples collected along the transect were take n from high elevations where possible (i.e.,

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96 closest to the original detachme nt level as possible) to ensu re that cooling ages obtained from the samples would reflect tectonic unr oofing and subsequent cooling of the ACC lower plate. Individual samples along the tr ansect were chosen based on their mineral content; samples with multiple argon-bear ing minerals (e.g., hornblende, K-feldspar, muscovite, and biotite) were targeted. Rock types sampled along the transect included: micaeous quartzites correlated with the lower Belt Supergroup and middle Cambrian section, undeformed granitoids, and deform ed granitoids with greenschist facies mylonitic fabrics from within the mylonite zone in the easternmost part of the study area. Four other rock samples were collected from areas outside the transect. Two granitoid samples were collected from expos ures of the detachment north of the study area, along the eastern flanks of the Flint Creek Range. The other two samples were collected from the detached upper plate with in the Deerlodge Valley, east of the study area; these include a large granodiorite block or mega-clast and a crystal-lithic rhyolitic tuff unit correlated with the Eocene Lo wland Creek Volcanic Sequence (Isopolatov, 1997). Previous Thermochronology Previously available thermochronological data relevant to the exhumation and cooling history of the ACC have been compiled and are summarized in Table 6-1. The majority of the previously thermochr onology was obtained using the K-Ar method. However, cooling ages reported in some of the more recent studies were obtained by 40Ar/39Ar and fission-track thermochronology. Th e relevant previous thermochronology are grouped by geographic location in Tabl e 6-1 (e.g., by name of mountain range or valley) and discussed briefly below. In addition, the previous K-Ar, 40Ar/39Ar and

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97 fission-track thermochronology summarized in Table 1 are presented on a regional scale geologic map in Figure 6-1. Four previous studies provide thermochr onological data from the Anaconda-Pintlar Range. Desmarais (1983) reports K-Ar cooli ng ages from granitoids of the Chief Joseph Batholith in the southwestern Anaconda-Pintl ar Range. Quartz di orite and granodiorite rocks, some exhibiting strong solid-state deformation, yield hor nblende and biotite cooling ages ranging from ~75-60 Ma. Biotite from a biotite granite a nd dacite dike yield K-Ar cooling ages of 51.3 1.6 and 50.3 2.2 Ma respectively; these two samples were collected east of the Chief Joseph Batholit h. Foster and Raza (2002) report apatite fission-track ages of ~40-30 Ma from granodi orite and quartz diorit e rocks of the Chief Joseph Batholith. Wallace et al. (1992) repor t K-Ar cooling ages from the central and northeastern Anaconda-Pintlar Range; biotite an d muscovite from a number of granodiorite samples and a dacite dike within in the central Anaconda -Pintlar Range give cooling ages of ~55-49 Ma. In addition, bi otite and hornblende from the Storm Lake Stock granodiorite (sample E1048) within th e northeastern Anaconda-Pintlar Range (and within the current study area) yield c ooling ages of 78.7 1.6 and 116.4 4.6 Ma, respectively. ONeill et al. (2004) report a 40Ar/39Ar cooling age of 47.2 0.3 Ma for muscovite from a mylonitic micaeous quartzite (sample ME-1) of the Sullivan Creek drainage in the northeastern Anaconda Pintlar Range (also within the current study area). Note rock samples were taken from the same localities as samples E1048 and ME1 for 40Ar/39Ar thermochronology in this study. Three previous studies pr ovide thermochronology from the Flint Creek Range. Hyndman (1972) and Martin et al. (1989) report hornblende and biotite K-Ar cooling

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98Table 6-1. Summary of releva nt previous thermochronology. Sample Rock type Latitude Longitude Mineral Age (Ma) Error (+/-)MethodSource Northern Flint Creek Range, Royal Stock (1) FC-263 porphyritic granite 464'30'' 11305'40'' bt64.7 1.6K-ArMartin et al., 1989 Southern Flint Creek Range, Mt. Powell Batholith (2) FC-273 granodiorite 468'23'' 11300'16'' ms61.5 1.0K-ArMartin et al., 1989 Western Flint Creek Range, Phillipsburg Batholith (3) BB6 hb bt granodiorite 466'19'' 11314'07'' hbl76.7 2.5K-ArHyndman et al., 1972 bt74.0 2.1K-Ar AA-4 hb bt granodiorite 468'21'' 11313'39'' hbl72.0 2.5K-Ar bt73.4 2.1K-Ar FC-271 hbl bt granodiorite 468'52'' 113 15'00'' bt77.2 1.9K-ArMartin et al., 1989 FC-272 hbl bt granodiorite 466'42'' 11311'17'' bt73.3 1.8K-Ar 98-75 hb bt granodiorite 469.1' 11313.45' ap57.0 5.0fis-trFoster and Raza, 2002 Northern Anaconda-Pintlar Range (4) E0134 granodiorite 46.55' 11310.51' ms 50.6 1.0K-ArWallace et al., 1992 bt48.5 1.0K-Ar E1048 hbl bt granodiorite 46.10' 11316.55' hbl116.4 4.6K-Ar bt78.7 1.6K-Ar bt53.1 1.6K-Ar ME-1 mylonitic mica quartzite 461'16" 113'04" ms47.2 0.3Ar-ArO'Neill et al., 2004 Sapphire Batholith (5) Grouped granodiorite, granite hb, bt~75-73 Ar-ArFoster, unpub. data 98-65 granodiorite, granite 463.44' 113 43.58' ap44.0 3.0fis-trFoster and Raza, 2002 98-66 granodiorite, granite 46 13.64' 11341.83' ap45.0 4.0fis-tr

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99Table 6-1 Continued. Sample Rock type Latitude Longitude Mineral Age (Ma) Error (+/-)MethodSource Chief Joseph Batholith (6) ND-80-173 hbl bt granodiorite 461'20'' 11349'10'' hbl74.9 2.6K-ArDesmarais, 1983 bt71.3 5.1K-Ar ND-80-298 q dioritef 456'20'' 11343'30'' hbl73.4 3.2K-Ar bt64.6 2.8K-Ar ND-81-475 q dioritef hb72.8 2.5K-Ar bt64.1 2.2K-Ar ND-80-14 hbl bt granodioritef 453'58'' 11348'56'' hbl66.8 2.9K-Ar bt61.8 1.9K-Ar AP-81-529 bt granodioritef 451'19'' 11349'20'' bt60.2 1.9K-Ar Grouped granodioritef see map ap~40-30 fis-trFoster and Raza, 2002 Central and southern Anaconda-Pintlar Range (7) ND-80-120 bt granodiorite 453'45'' 11354'00'' bt57.6 2.5K-ArDesmarais, 1983 ND-81-362 biotite granitef 454'30'' 11353'40'' bt59.7 2.6K-Ar ND-79-269 two-mica granodiorite 451'35'' 11350'58'' ms58.2 2.1K-Ar bt54.8 2.4K-Ar ND-79-89 bt granite 456'25'' 11357'25'' bt51.3 1.6K-Ar AP-81-123 dacitic dike 450'33'' 11342'53'' bt50.3 2.2K-Ar 81642 granodiorite 456.04' 11328.40' ms49.9 1.0K-ArWallace et al., 1992 bt50.7 1.0K-Ar

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100Table 6-1 Continued. Sample Rock type Latitude Longitude Mineral Age (Ma) Error (+/-)MethodSource 81540 granodiorite ms54.9 1.2K-Ar 81639 granodiorite 454.21' 11331.14' bt53.1 1.2K-Ar 81640 dacitic dike 455.30' 11330.17' bt51.8 1.0K-Ar Pioneer Batholith (8) Grouped q diorite, granodiorite, see map hb, bt, ms~80-65 K-Ar, Ar-ArSnee, 1978, 1982 granite Boulder Batholith (9) Grouped granodiorite, granite see map hb~76-71 K-ArTilling et al., 1968 bt~74-70 K-Ar Grouped late-stage granitic veins see map kfs, bt, ms~74-59 Ar-ArLund et al., 2002 Deerlodge Valley (10) LVC-32 rhyolite 468'00'' 11256'2 0'' sd48.4 0.5Ar-ArIsopolatov, 1997 LVC-6 rhyodacite porphyry 461'10'' 11258'50'' hbl51.7 1.1Ar-Ar bt51.5 0.4Ar-Ar LCV-19-3 andesite porphyry 463'10'' 11244'30'' plTF52.6 1.9Ar-Ar LCV-15 Dacitic porphyry pl49.8 0.5Ar-Ar LCV-18 andesite porphyry 463'10'' 11244'30'' plTF50.2 2.1Ar-Ar

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101Table 6-1 Continued. Sample Rock type Latitude Longitude Mineral Age (Ma) Error (+/-)MethodSource 95LCV-9 rhyodacite porphyry 460'47'' 11244'37'' hbl52.2 0.4Ar-Ar 95LCV-7 rhyolite porphyry 460'57'' 11242'11'' hbl52.6 0.2Ar-Ar 95LCV-10A rhyolite tuff 460'20'' 11226'44'' bti52.4 0.5Ar-Ar 00-11 rhyolite tuff bt52.9 0.4Ar-ArDudas, unpub. data 00-24 rhyolite tuff bt51.5 0.4Ar-Ar BQM-2 rhyodacite porphyry bt51.5 0.2Ar-Ar Elkhorn Mts. Volcanic Fie (11) Grouped varied hbl, bt~81-74 K-ArTilling, et al., 1968 Note: q = quartz, bt = biotite, ms = musc ovite, hbl= hornblende, kfs = K-feldspar, sd = sanidine, pl = plagioclase, ap = apatit e, = sample also dated in this study, f = moderate to strong solid-state foliation, TF = total fusion age, i = normal isochron age, MSWDp = mean standard weighted deviates for plateau cooling age, PB = Philipsburg Batholith, MPB = Mt. Powell Batholith, RS = Royal Stock, C J = Chief Joseph Batholith. Number geographic regions correspond to Figure 6.1. All cooling ages are reported with 2 sigma error.

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102 Figure 6-1. Geologic map of Anaconda me tamorphic core complex (ACC), western Montana and vicinity showing locati on of previous thermochronology. See Table 1 for explanation for number regions. RS = Royal Stock, Pb = Philipsburg Batholith, MPb = Mt. Powell Batholith, SLP = Storm Lake Stock pluton, SP = Sapphire Batholith, CJ = Ch ief Joseph Batholith, PB = Pioneer Batholith, BB = Boulder Ba tholith, LCV = Lowland Creek Volcanic Field, EHV = Elkhorn Mountains Volcanic Field. The current study area is designated by the dashed box (Modifi ed from Foster et al., 2006a). ages ranging from ~77-72 Ma for the Philipsburg Batholith, a predominately granodioritic intrusion in th e northwestern Flint Creek Range In addition, Foster and

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103 Raza (2002) report an apatite fission-track c ooling age of 57.0 5 Ma for a sample taken from the western part of the Philipsburg Batholith. Marvin et al. (1989) pr ovide biotite and muscovite K-Ar cooling ages of ~65-60 Ma for the Royal Stock (mostly granodiorite) and Mount Powell Batholith (g ranodiorite and two-mica granite) in the northeastern Flint Creek Range. Isopolatov (1997) and Dudas (unpublished data) report 40Ar/39Ar cooling ages from several samples collected from the Lowland Creek Volcanic Field (LCV) located within the Deerlodge Valley, east of the Anaconda-Pin tlar and Flint Creek Ranges. Note some of these samples were collected from the we stern Deerlodge Valley, di rectly east of the current study area within the de tached, brittlely fau lted upper plate of the ACC. Biotite, hornblende, plagioclase, and sa nidine from rhyolitic lava flows/tuffs and rhyolitic-toandesitic hypabyssal (very shallow) porphyries of the LCV all report 40Ar/39Ar cooling ages ranging from ~53-48 Ma. Note that a sample (DF02-113) was collected from the same outcrop as LCV sample LV-6 (of Isopolatov, 1997) for 40Ar/39Ar thermochronology in this study (Table 6-1). The cooling histories of the Pioneer, Boul der, and Sapphire Batholiths also bear significant importance on later discussions of the exhumation and cooling history of the ACC. Therefore, a summary of the thermochronology from these batholiths is also included here. Snee (1978, 1982) provides a large K-Ar and 40Ar/39Ar thermochronological data set for the Pioneer Batholith, a predominately quartz dioritegranodiorite-granite intrusion located south east of the Anaconda-Pintlar Range and east of the Bighole Valley. Hornblende, biotite a nd muscovite cooling ages from the Pioneer Batholith reported by Snee range from ~80-65 Ma. The Boulder Batholith is a large

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104 (>6000 km2) intrusion, mostly comprised of a gran ite phase (previously classified as a quartz monzonite; see Lund et al., 2002) and se veral smaller granodior ite intrusions that lies directly east of the current study area a nd east of the Deerlodge Valley (Vejmelek and Smithson, 1995). Tilling et al. (1968) repor t biotite and hornblende K-Ar cooling ages of ~76-71 Ma for plutons of the Boul der Batholith and ~81-74 Ma for the Elkhorn Mountains Volcanic Field, the volcanic cove r for the Boulder Batholith. In addition, Lund et al. (2002) report 40Ar/39Ar cooling ages of ~74-59 Ma for late-stage granitic veins intruded within the main phases of the Boulde r Batholith. Foster (unpublished data) reports muscovite and biotite 40Ar/39Ar cooling ages of ~75-73 Ma from the Sapphire Batholith located west of the cu rrent study area and eas t of the Bitterroot Valley. Foster and Raza (2002) also report apatite fission-tr ack cooling ages of 44.0 3 and 45.0 4 Ma for two samples collect ed from the Sapphire Batholith. 40Ar/39Ar Thermochronology Results Individual mineral separate s (biotite, muscovite, hornble nde, and K-feldspar) were obtained from a total of the twenty-two ro ck samples (twenty lower plate and two upper plate samples) for the 40Ar/39Ar thermochronology in this study. These mineral separates were subsequently analyzed during in-vacuo laser and furn ace step-heating experiments coupled with mass spectrometry in the noble gas laboratory at the University of Florida. Sample preparation, analytical instrumenta tion and procedures, and thermochronological data reduction are summarized in Appendix A. The results from the 40Ar/39Ar thermochronology of this study are reported in Table 6-2. These results are organized into tw o groups; one corresponds to rock samples collected along the lower plat e transect described above a nd the other group to samples collected outside the lower plate transect. Note the results from rock samples collected

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105 along the lower plate transect are listed spatially, from west to east (top to bottom in Table 6-2) according to their relative position along the lower plate transect. Individual 40Ar/39Ar cooling ages are reported in Table 62 according to rock sample name, rock type, sample location, elevation, an d mineral dated. Total percent 39Ar gas released during step-heating and MSWD (mean standard weighted deviates) are also reported with weighted plateau cooling ages in Table 6-2. The J-values (from mineral flux monitors) used in individual cooling age calculations ar e also included in Table 6-2. In addition, mineral cooling ages obtained from samples collected along the lower plate transect are shown with sample localities on a simplified geologic sketch map in Figure 6-2. The individual mineral cooling age sp ectra are shown in Figure 6-3. Argon Closure Temperature Mineral ages obtained using 40Ar/39Ar thermochronology in this study are cooling ages, not crystallization ages. Therefore, th ese cooling ages record the time at which a mineral grain (and its host rock) cooled below a certain temperature and when 40Ar, the radiogenic daughter product of 40K, fully ceased to escape (diffuse) from the mineral crystal lattice. This critical temperature is referred to as the closure temperature, below which the mineral system is effectively cl osed to the diffusion of Ar (Dodson, 1973; Harrison and McDougall, 1999). Ar closure temperatures vary between different minerals in a given rock because of significan t differences in their crystalline structures and consequently, in Ar diffu sion properties. Modeling a nd experimentation involving Ar diffusion over the past several decade s has provided approximate Ar closure temperatures for common K-bearing minera ls are: ~300C for biotite, ~350C for muscovite, and ~500C for hornblende (see Harrison and McDougall, 1999 and many references therein). K-feldspar has been shown to have a range of Ar closure.

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106Table 6-2. Summary of 40Ar/39Ar thermochronology from the Anaconda metamorphic core complex. Sample Rock type Latitude Longitude Elevation (m) Mineral Age (Ma)Errors (+/-)% 39ArpMSWDp Samples from the 40Ar/39Ar transect in the northeastern Anaconda-Pintlar Range SL-38 hbl bt grd 46.10' 11316.55' 2311 bt 79.01.2782.02 hbl U102.512.865144.50 kfs LT56.91.694.98 WG04-114 hbl bt grd 464'25'' 11315'48' 2503 bt 74.00.9940.69 WG04-112 micaeous quartzite 463'26" 11314'43" 2646 ms 56.90.7810.34 bt 52.80.71000.63 WG04-052 q di 465'09" 11313'46" 2754 bt 64.11.0881.43 WG04-033 dacite dike 464'29" 11310'50" 2583 bt 48.50.6940.99 WG04-109 two mica gr 462'35" 11310'43" 2878 ms 48.00.6871.75 bt 47.60.61001.60 WG04-101 coarse two mica gr 463'39" 11310'01" 3073 ms 49.60.6780.34 bt 49.20.9931.25 WG04-100 pegmatitic dike 463'36" 11309'52" 2927 ms 50.40.8952.63 ME-6 coarse two mica gr 462'25" 11310'15" 2732 ms 51.51.2790.31 bt 48.61.0951.74 Ug-1 deformed q di sill 464'32" 113'55" 2610 bt 50.50.8920.10 hb U91.31.0600.61 WG04-103 mylonitic quartzite 461'16" 11310'04" 2524 ms EP47.31.1800.91 WG04-092 bt gr 463'34" 11308'25" 2896 bt 47.10.7783.01 WG04-089 aluminous leuco-gr 462'56" 11308'18" 2646 ms 45.50.61001.43 WG04-138 mylonitic quartzite 46 3'37" 11306'14" 3098 ms 46.10.6990.19 DF02-116a mylonitic two mica gr 465'29" 11302'05" 2573 ms 40.52.0652.21 bt 39.62.3701.36 bt TF40.70.7100DF02-120 mylonitic biotite gr 463'25" 11301'38" 2229 bt TF38.50.7100

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107Table 6-2. Continued. Sample J value Comments SL-38 0.0038300 0.0038500 extraneous argon 0.0038450 WG04-114 0.0005150 WG04-112 0.0005140 0.0005140 WG04-052 0.0005145 WG04-033 0.0005150 WG04-109 0.0005140 0.0005140 WG04-101 0.0005150 0.0005150 WG04-100 0.0005150 ME-6 0.0064000 0.0064000 Ug-1 0.0038590 0.0038500 extraneous argon WG04-103 0.0005145 WG04-092 0.0005150 WG04-089 0.0005140 thin garnet bearing sheet WG04-138 0.0005150 Short Peak mylonite DF02-116a 0.0038000 Clear Creek mylonite 0.0037650 0.0064600 DF02-120 0.0064770 Mill Creek mylonite

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108Table 6-2. Continued. Sample Rock type Latitude Longitude Elevation (m) Mineral Age (Ma)Errors (+/-)% 39ArpMSWDp Outside the 40Ar/39Ar transect DF02-118a mylonitic bt gr 46 1'49" 11259'11" bt 38.81.6720.82 DF02-118b mylonitic bt g 461'49" 11259'11" bt 42.41.61001.03 DF02-121 gr bt 55.32.8830.84 DF02-119a two mica gr 466'42" 11257'53" ms EP68.32.3984.14 bt EP66.21.2812.23 DF02-113 rhyolite tuff 461'11" 11259'03" bt 53.71.4660.71 DF02-114 bt hbl gr 464'24" 11256'22" bt 76.31.1950.57 Table 6-2. Continued. Sample J value Comments DF02-118a 0.0038100 Lost Creek mylonite DF02-118b 0.0038540 Lost Creek mylonite DF02-121 0.0038500 DF02-119a 0.0038540 Racetrack Creek 0.0038540 DF02-113 0.0037650 upper plate, lower LCV DF02-114 0.0038770 from upper plate Note: grd = granodiorite, q di = quartz diorite, gr = granite, q = quartz, bt = biotite, ms = muscovite, hbl = hornblende, kfs = K-feldspar, = previously dated by other workers, LT = low temperature plate au, TF = single grain total fusion cooling age, EP = error plateau u = unreliable cooling age, % 39Arp = percent of 39Ar used in weighted platuea age calcuation, and MSWDp = mean standard weighted deviates for plateau cooling age. All cooling age are reported in 2 sigma error.

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109 Figure 6-2. Geologic sketch map showing mineral cooling ages from samples coll ected along the ACC lower plate transect using 40Ar39Ar thermochronology. Mineral cool ing age errors are two sigma.

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Figure 6-3. 40Ar/39Ar mineral age spectra obtained for samples collected from the Anaconda metamorphic core complex using 40Ar/39Ar thermochronology. All ages are weighted plateau cooling ag es calculated from three or more contiguous heating step s with >50% total 39Ar gas released unless otherwise specified. Error plateau ages were calculated from noncontiguous heating steps and/or less than 50% total 39Ar gas released. All ages are reported with two sigma errors. Thickness of age plateaus corresponds to the two sigma errors for the cooling age calculated fo r that particular heating step.

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111 Figure 6-3. Continued.

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112 Figure 6-3. Continued

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113 Figure 6-3. Continued.

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114 Figure 6-3. Continued.

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115 Figure 6-3. Continued.

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116 Figure 6-3. Continued.

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117 Figure 6-3. Continued.

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118 Figure 6-3. Continued.

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119 Figure 6-3. Continued.

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120 Figure 6-3. Continued. temperatures from ~200-250C (lower temper ature closure) to ~350C (high temperature closure); the variation in Ar cl osure temperature for K-feldspar arise from the presence of multiple sized Ar diffusion domains within its crystalline structure (e.g., Lovera et al., 1989; Foster et. al., 1990). Because of the differences in the Ar closure temperature between K-bearing minerals cooling histories can be derived fr om single rock samples containing more than one of these minerals. Note, however, that Ar closure temperature for any given Kbearing mineral is a function of lithosta tic pressure, Ar diffusion domain size, and

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121 especially cooling rate (dT/dt) (Lister and Baldwin, 1996). Therefore, these factors must be taken into account when estimating Ar cl osure temperatures and using cooling ages obtained from K-Ar or 40Ar/39Ar thermochronology to constrai n the cooling histories of exhumed rocks. In this study, the Ar closure temperatures (Tc) for biotite, muscovite, and hornblende were calculated using the follo wing formula from Harrison and McDougall (1999): Tc = (E/R)/ln[(ARTc 2(Do/r2))/(E(dT/dt))] (6-1) where E is the diffusion activation energy, R is the gas constant (1.987 cal/mol-K), A is the diffusion geometry coefficient, Do is the frequency factor, r2 is the effective diffusion domain radius, and dT/dt is the estimated cooling rate. The diffusion parameters for muscovite and hornblende used in equation X were taken from Lister and Baldwin (1996). The diffusion parameters for biotite are from Harrison et al. (1985). The diffusion parameters are summarized in Table 6-3. The cooling rates for both biotite and muscovite (dT/dt) are estimated to be 100C/Ma because biotite and muscovite cooling ages from two two-mica granite samples in the central study area (WG04-101 and WG04-109) are essentially equal (well within analyt ical error, Table 3). Therefore, a cooling rate of 100C/Ma was used in the biotite and muscovite closure temperature calculations here. A cooling ra te of 100C/Ma was also applie d to the hornblende closure temperature calculati on. For biotite Tc = 343 20C, for muscovite Tc = 390 20C, and for hornblende Tc = 545 20C. The 20C error wa s applied to these calculations to estimate uncertainties in the diffusion para meters used and in Ar closure temperatures at pressures of 0-5 kbar (see Fig. 2a, p. 88, Lister and Baldwin, 1996). The closure temperature of K-feldspar was not calculated here and is taken to be ~200-250C for

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122 lower temperature closure and ~ 350C for hi gh temperature Ar closure. Below the results of the 40Ar/39Ar thermochronology from this study are summarized by mineral type. When a cooling age is reported here for a given mine ral the age corresponds to the Ar closure temperatures calculated or estimated here. Biotite Laser step-heating experiments and data re duction from eleven biotite separates across the lower plate 40Ar/39Ar transect resulted in mostly well-defined, flat-shaped weighted age plateaus (Figure 6-3).1 In every experiment, incremental heating steps were highly radiogenic (i.e., very low 36Ar/40Ar ratios). In most cases the highly radiogenic steps clustered low on the inverse isochrons but within error of atmospheric Ar (36Ar/40Ari = 3.38 x 10-3) Therefore, regression of these heating steps produced unrealistic regressions on inverse isochron di agrams, hence poorly defining the initial values of 36Ar/40Ar in the biotites. Nonetheless, in all but one of the analyses the inverse isochron ages are within error of the weighted plateau ages (Table X). This is because most incremental heating steps used in the age calculations yielded 36Ar/40Ari ratios within error or very close to atmospheric 36Ar/40Ari. However, the weighted plateau ages calculated from the heating experiments are c onsidered the most reliable and preferred over inverse isochron ages. Apparent biotite weighted plateau ages de fine a lateral age gradient across the ACC lower plate, where biotite cooling ages young in a ESE direction along the 40Ar/39Ar transect (Table 2, Fig. 3). In the westernm ost transect biotite from the SLS yielded 1 The definition or criterion used to define a 40Ar/39Ar age plateau is not uniquely defined in the literature (e.g., see McDougall and Harrison, 1999). However, in this study a reliable age plateau is considered to be any age plateau consisting of 50 total percent 39Ar released from three or mo re contiguous heating steps; the contiguous steps must be within 2 error as well.

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123 weighted plateau ages ranging from 79-64 Ma Biotite from sample SL-38 (2311 m), a granodiorite sampled from the SLS, yielded a flat age plateau over 78 percent of the total 39Ar gas released corresponding to a coo ling age of 79.0 1.1 Ma (MSWD = 2.02). The higher MSWD for this age calculation is due to the presence of a small amount of excess Ar in the first and last few heating steps of the experiment. However, the plateau age is considered robust with these anomalous step s omitted. Notably, the biotite cooling age from SL-38 is concordant with a K-Ar bi otite cooling age of 78.7 1.6 Ma from E1O48, a sample collected from the same outcrop by Wallace et al. (1992, see Table 1). Biotite from WG04-114, another granodi orite sampled from the SLS near the eastern shore of Storm Lake (2503 m) gave a plateau age of 74.0 0.9 Ma (94% total 39Ar gas, MSWD = 0.69). In addition, biotite from sample WG04-052, a quartz diorite from the SLS, yielded a substantially younger weighted pl ateau age of 64.1 1 Ma (88 % 39Ar gas, MSWD = 1.43). This sample was collected from a ri dge top (2754 m) just east of and overlooking the Fourmile Lakes Basin along the western pa rt of the transect (Appendix D). Biotite cooling ages abruptly young east and south of the SLS. Biotite from WG04-112, a micaeous quartzite gneiss sampled south of the SLS in the upper Twin Lakes Creek drainage yielded a flat age plateau over 100 percent 39Ar gas release and a cooling age of 52.8 0.7 Ma (MSWD = 0.63). East of Lake of the Is le, in the center of the transect, biotite from two samples yield weighted plateau ages near 50 Ma. Biotite from a highly deformed hornblende-biotite quartz diorite sill (Ug-1) ga ve a plateau age of 50.5 0.8 Ma over 92 percent total 39Ar gas released (MSWD = 0.10). Biotite from an undeformed coarse-grained two-mica granite (WG04-101) sampled along the continental divide (3073 m) yielded an ag e plateau showing slight Ar lo ss over the first two heating

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124 steps. However, with these steps excluded a cooling age of 49.2 0.9 Ma was calculated over the remaining 93 percent of the total 39Ar gas released (MSWD = 1.25). An aphanitic porphyritic dacite dike sampled near the Lake of the Isle yielded a slightly younger biotite plateau age of 48.5 0.6 over 94 percent of its 39Ar gas (MSWD = 0.99). This dacite dike crosscuts the foliated quartz diorite sill and highly deformed metasedimentary strata in central LISZ. In addition, the dacite dike and others like it crosscut an undated medium-grained two-mica granite pluton south of the Lake of the Isle (Appendix D). Two samples of undeformed coarse-grain ed two-mica granite (similar to WG04101) collected south of the central transect on the southeastern flank of Mt. Evans yielded biotite cooling ages of ~ 48 Ma. WG04109 (2878 m) gave a cooling age of 47.6 0.6 Ma over 100 percent of its 39Ar gas release (MSWD = 1.6). Biotite from ME-6 sampled nearby but down slope (2732 m) of WG04-109 yi elded a flat age plateau over 95 percent the total 39Ar gas released and a similar cooling age of 48.6 1.0 Ma. Wallace et al. (1992) reports a K-Ar biotite cooling age of 48.5 1 for two-mica granite sampled from the same outcrop as WG04-109. Eastward along the transect an undeforme d biotite granodiorite (WG04-092) was sampled from a steep cliff just below the c ontinental divide (2896 m) on the eastern side of Tenmile Lakes Basin. Biotite from this sample yielded a plateau indicative of some mild Ar loss over the first 2 steps. Howe ver, a cooling age of 47.1 0.7 (MSWD = 3.01) was calculated from the remaining 78 percent 39Ar gas of the experiment defining a flat plateau.

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125 Step heating experiments carried out on biotite from the greenschist facies twomica granite mylonite (DF02-116a) from uppe r Clear Creek yielded a U-shaped age plateau indicating the presence of some excess Ar. However, three contiguous steps are within two sigma error and yielded a we ighted plateau age of 39.6 2.3 Ma over 70 percent of the total 39Ar gas release (MSWD = 1.36). A total fusion age of 41.7 3.2 Ma was calculated from the same analysis. The la rger error associated with the plateau age for DF02-116a is due to low Ar yields fr om the incremental h eating steps (0.002-0.4 V 40Ar). Note a separate total fusion analys is of a single biotit e grain from DF02-116a yielded a cooling age of 40.7 0.7 Ma, within error of the weighted plateau and total fusion ages from the first analysis. These data indicate that DF02-116a cooled below the biotite closure temperate by about 40 Ma. No step-heating experiments conducted on DF02-120, the greenschist facies biotite granit e mylonite from Mill Creek. However, a total fusion analysis of a singl e biotite grain from this samp le yielded a cooling age of 38.6 0.7 Ma, slightly younger than the biotit e total fusion age from DF02-116a. Muscovite Muscovite separates from nine different rock samples collected from the ACC lower plate underwent in-v acuo laser step-heating 40Ar/39Ar analyses. Note five of these muscovite analyses are paired wi th biotite analyses from the same samples (Table 2). As in the biotites, the analyses of muscovite resulted in largely well-defined, flat-shaped age spectra with highly radiogenic incrementa l heating steps. Thus, inverse isochrons from the muscovite Ar data are also ch aracterized by clustered data points near atmospheric Ar, poor 36Ar/40Ari regressions and inadequate 36Ar/40Ari estimations. Therefore, the weighed plateau ages for mu scovites are considered most reliable and preferred over invers e isochron ages.

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126 Apparent muscovite plateau ages genera lly young in an ESE direction across the ACC lower plate, similar to the lateral age pa ttern revealed from biotite cooling ages. In the western transect muscovite from WG 04-112 gave a plateau age of 56.9 0.7 over 81 percent of the total 39Ar gas released (MSWD = 0.32). Muscovite of sample WG04-101 from the central transect yielded a flat age plateau over 78 percent of the total 39Ar gas released and a cooling age of 49.6 0.6 Ma (MSWD =0.34). In addition, muscovite from WG04-100 (2927 m), a pegmatitic dike sampled nearby WG04-101 yielded an indistinguishable plateau age of 50.4 0.8 Ma (95 % 39Ar gas, MSWD = 2.63). South of the central transect, muscovite from samples WG04-109 and ME-6 gave plateau ages ranging from ~ 48-51 Ma. WG 04-109 yielded a well-constrained muscovite plateau age of 48.0 0.64 (87% total 39Ar gas, MSWD = 1.75). However, muscovite from ME-6 collected less than 1 km to the southeast of the WG04-109 sample locality gave a slightly older weighted plateau age of 51.5 1.2 Ma over 79 % of the total 39Ar gas released (MSWD = 0.31). Note the ag e plateau for muscovite from ME-6 is Ushaped indicating the presence of some excess Ar in the first and last few heating steps of the experiment. However, the plateau age quot ed here is reliable when the six heating steps with anomalous Ar are excluded. An i nverse isochron from the analyses of the ME6 muscovite shows that the five remaining contiguous heating steps used in the age calculation lie on a well-defined regr ession within error of atmospheric 36Ar/40Ari (Fig X). Muscovite from WG04-103 (2524 m), a mu scovite-bearing mylonitic quartzite, gave a weighted plateau age of 48.1 0.6 Ma over 80% 39Ar gas release (MSWD = 0.70). This sample was collected from a large mylonitic quartzite outcrop along the

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127 eastern side of the Sullivan Creek drainage, 2 km south of the ME-6 sample locality (Fig X). ONeill et al. (2004) report a sligh tly younger muscovite plateau age of 47.2 0.3 Ma (quoted as 2 error here) from a sample (ME-1) collected from the same locality (85 % 39Ar gas release, no MSWD reported). The s light difference in th ese cooling ages can be explained by a small amount of excess Ar in the WG04-103 muscovite as indicated by the inverse isochron, where three heating st eps used in the age calculation fall below atmospheric 36Ar/40Ari (i.e., have slight excess 40Ar). However, because the data points are clustered on the inverse isochron diag ram the excess Ar cannot be accurately quantified and subsequently corrected for. A thin (< 0.5 m) garnet-bearing leucogran ite vein was sampled from the biotite granodiorite in the upper Tenmile Creek draina ge along the eastern transect. The vein trends NE, roughly parallel to the nearby c ontact between the biotite granodiorite and porphyritic two-mica granite (Appendix D). In some places the leucogranite vein strikes parallel to a 1-2 m thick foliated zone in the biotite granodiorite. However, the vein can also be seen crosscutting this foliation; th erefore it post-dates emplacement of the biotite granodiorite. Analyses of muscovite from sample WG04-089 (2646 m) gave a very flat age plateau corresponding to cooling age of 45.5 0.6 Ma ove r 100 percent of the total 39Ar gas released (MSWD = 1.43). Notably, this cooling age a pparently ~ 1.6 m.y. younger than the biotite cooli ng age obtained from WG04-092 lo cated 1 km to the north and ~ 4-6 m.y. younger than muscovite cooli ng ages obtained from the porphyritic twomica granite just ~ 2.5 km to the west. WG04-138 was collected from a mylonitic muscovite-bearing quartzite unit near the summit of Short Peak (3098 m) on the con tinental divide between the Tenmile Creek

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128 and Mill Creek drainages. The mylonitic qua rtzite that outcrops at Short Peak is essentially identical to the mylonitic quartz ite exposed in the Sullivan Creek drainage (i.e., WG04-103). Muscovite from WG04-138 yielded a cooling age of 46.1 0.6 Ma over five contiguous heating step s and 99 percent of the total 39Ar gas released (MSWD = 0.19). As in the biotites, muscovite from the easternmost transect gave the youngest cooling ages. Muscovite from DF02-116a, th e Clear Creek two-mica granite mylonite, yielded a slightly U-shaped age plateau i ndicating some excess Ar in the sample. However, three contiguous heating steps from a flat part of the U-trough lie within 2 error and were used to calculate a we ighted plateau age of 40.5 2.0 Ma (65% 39Ar gas, MSWD = 2.21). The inverse isochron diagram pl otted from these data indicates the three heating steps used in the age calcula tion are within error of atmospheric 36Ar/40Ari, therefore excluding the excess 40Ar (Fig X). The cooling age of DF02-116a indicates that the Clear Creek two-mica granite mylonite c ooled below the muscovite closure by about 40 Ma. Hornblende Hornblende separates from two samples were analyzed by furnace step-heating to provide higher temperature age constraints for parts of the ACC lower plate. In particular, hornblende from SL-38 was analyzed to better constrai n the cooling history for the SLS granodiorite in the westernmost transect. In addition, hornblende from the sample Ug-1, the deformed quartz diorit e sill, was analyzed to provide higher temperature age constraints in the central tr ansect. A hornblende cooling age from this sill could be used to further constrain the c ooling history of the LISZ and to assess the

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129 extent of heating from mylonitization a nd emplacement of younger granitoids near this part of the lower plate. Unlike biotite and muscovite, step-heati ng experiments performed on these two hornblende separates resulted discordant and saddle-shaped age spectra. The age spectrum for the SL-38 hornblende is highly di scordant, where the firs t four heating steps of the experiment alternate between ~ 110 a nd ~ 90 Ma then ages steadily climb to ~ 183 Ma at the last step. This type of step-heating pattern in hornblendes is attributed abundant excess Ar in the sample (e.g., Rich ards and McDougall, 1990). K/Ca ratios from the sample anomalously high (0.8-1.1) over the first two heating steps of the experiment (800 and 900C) suggesting th e SL-38 hornblende contains K-bearing mineral inclusions. Such inclusions could account for the excess Ar or inherited Ar (Kelley, 2002) in the first tw o heating steps of the experiment. However, K-bearing mineral inclusions can not account for the ap parent excess Ar in the remaining heating steps because all give very low K/Ca ratios (<0.1). Despite the obvious presence of excess Ar in the SL-38 hornblende some useful information can be extracted from the result s of this experiment. The SL-38 hornblende yielded a total fusion age of 108.8 1.7 Ma, ab out 30 Ma older than the biotite cooling age from the same sample. Wallace et al (1992) report a similar spread between hornblende and biotite K-Ar cooling ages for the SLS granodiorite sampled from the same locality (~ 37 Ma difference). The appa rent difference is these cooling ages imply slow cooling of SLS granodiorite, on the or der of < 10C/Ma. However, removal of the now documented excess Ar in SLS hornble nde would result in a more realistic hornblende cooling age and a more rapid coolin g rate. Regression of four heating steps

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130 on an inverse isochron diagram can roughly quan tify some of the exce ss Ar in the SL-38 hornblende (36Ar/40Ari of 1165 140, Fig X). The inverse isochron age calculated from this regression is 80.5 7 Ma. This cooling ag e is within error of the biotite weighted plateau age of SL-38 suggesti ng rapid cooling (~ 200C/Ma) of the SLS granodiorite at ~ 79 Ma rather than slow cooling over 30 Ma. The Ug-1 hornblende yielded a strongly sa ddle-shaped age spectrum where the first and last few heating steps of the experiment gave anomalously old cooling ages (~ 330460 Ma). Ages from the steps between these ranged from ~ 180-70 Ma. This hornblende age spectrum is i ndicative of abundant excess Ar as well. However, unlike hornblende from SL-38, all heating steps from th is experiment yielded low K/Ca ratios (< 0.02). Thus, the apparent excess Ar in the Ug -1 hornblende is not likely derived from Kbearing mineral inclusions. One possible source for excess Ar in this sample is Ar uptake during a reheating event. Harrison and McDougall (1980) report age spectra with anomalously high ages in early heating steps from amphiboles sampled within the contact aureole adjacent of a Cretaceous intrusion. Th ey attribute the anomalously old ages in parts of the age spectra to ex cess Ar incorporated into gr ain margins during reheating by the neighboring intrusion. The Ug-1 sample was collected < 100 m fr om the contact with a younger two-mica granite intrusion (Fig X). Thus, reheating by the two-mica granite intrusion may have allowed excess Ar to be incorporated into the Ug-1 hornblende. The fact that biotite from Ug-1 reports an Eocene cooling age (50.5 0.8 Ma) suggests temperatures were at least high enough to practically reset Ar in biotite. Unfortunately, the excess Ar in the Ug-1 hornblende could not be quantified and corrected for using an inve rse isochron diagram. Most heating steps reported 36Ar/40Ar

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131 values 0, lying along the x-axis of the diagram. The low 36Ar/40Ari ratios in the Ug-1 hornblende are the result of very high radiogenic (high 40Ar) and low 36Ar yields over most of the heating step s. Thus, a realistic 36Ar/40Ari ratio needed to correct for excess Ar could not be obtained from the data. Th erefore, the total fusion age of 113.3 1.4 Ma from the analysis of Ug-1 hornblende is considered geologica lly unsound because of abundant uncharacterized excess Ar. Note als o, this total fusion age is ~ 38 Ma older than the U-Pb zircon crystallization ag e obtained for the same sample further exemplifying its unreliability (s ee U-Pb geochronology results). K-feldspar One K-feldspar separate from SL38 underwent furnace step-heating and 40Ar/36Ar analysis. Cooling age constraints from the SL38 K-feldspar were desired to be used in conjunction with the U-Pb zircon, hornblende a nd biotite data from the same sample to provide an in depth cooli ng history for the SLS granodi orite. Vacuo step-heating experiments of K-feldspars have shown K-feldspar is characterized by multiple Ar diffusion domains of different sizes, therefor e having different Ar closure temperatures (e.g., see Lovera et al., 1989). The Ar clos ure temperature for K-feldspars typically ranges from ~200-250C for the smaller di ffusion domains to ~ 350-400C in larger domains. Therefore, modeling of Ar data obtained from step-heating analyses of Kfeldspars can used to produce time-temperature curves that often span several million years (Lovera, 1992; Lovera et al., 2002). A single K-feldspar step-heating analysis and modeling can often reveal a de tailed cooling history for a given rock sample in itself (e.g., Richter et al., 1991). Furnace step-heating of the SL-38 K-feldspar produced an age spectrum characterized by anomalously old cooli ng ages over the first few percent 39Ar released

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132 then ages that rapidly drop to ~ 6555 Ma for the next 30% cumulative 39Ar release. The cooling ages then gradually c limbed to near 100 Ma with one anomalous step yielding a cooling age of ~ 118 Ma. Thereafter, cooling ag es abruptly rise to > 300 Ma over the last few percent cumulative 39Ar released (Fig. 3). The gra dual increase in ap parent cooling ages over much of cumulative 39Ar released is common in many step-heating experiments of K-feldspars. Such an age gradient refl ects release of Ar from diffusion domains with different Ar closure te mperatures with progressive heati ng of the sample (Lovera et al., 1989; Harrison et al., 1991). However, the anom alously old cooling ages of the first and last few percent cumulative 39Ar released from the SL-38 K-feldspar are unmistakably due to the presence of excess Ar (McD ougall and Harrison, 1999). Harrison and McDougall (1981) attribute such saddle-shaped K-feldspar age spectra to excess Ar trapped in cation cites and anion vacancies in the near-surface crystal lattice of Kfeldspar at temperature < 350C (see Foster et al., 1990 for a review). An inverse isochron diagram from the step -heating experiment of SL-38 K-feldspar shows that virtually all steps released have abundant excess Ar (i .e., all but two heats yielded 36Ar/40Ari ratios of < 3.38 x 10-3). In fact, the last 50% cumulative 39Ar release from the experiment yielded unreasonable ages, older than the U-Pb zircon crystallization age for WG04-114, another SLS granodiorite sa mple. Therefore, the age spectrum for the SL-38 K-feldspar is largely unreliabl e and unsuitable for detailed K-feldspar modeling. Thus, a detailed co oling curve is obtained fo r the SL-38 K-feldspar. However, some useful time-temperature constrains can be deduced from the lower temperature portion of this step-heating experiment.

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133 The heating steps over 2-20 percent cumulative 39Ar released during the step heating experiment define saw tooth patt ern where older cooling ages alternate with slightly younger cooling ages. These alte rnating steps correspond to the isothermal heating steps employed in the step-heating schedule of the SL-38 K-feldspar (550750C). Heating steps that yield slightly younger cooling ages are from the second fraction of Ar gas analyzed at each temperat ure, and vise versa. Harrison et al. (1994) demonstrated that the excess 40Ar in the older steps of this saw tooth pattern is strongly correlated with Cl probably derived from flui d inclusions in K-feld spars (i.e., a strong linear relationship between 40Ar/K and Cl/K ratios). This excess 40Ar is often released from fluid inclusions during low temperatur e steps of the experiment creating this common discordant saw tooth pattern. Therefore, by discarding the first of each isothermal heating step the Cl-correlated excess 40Ar is avoided and useful timetemperature constraints can be deduced from the low temperature (< 800C) portions of K-feldspar step-heating experiments. Followi ng this strategy, four Cl-corrected excess Ar steps were removed from eight steps in th e low temperature portion of the age spectra. The four remaining steps were used to calcu late an error plateau age of 56.9 1.61 Ma. Note the four steps used in this age calculation lie close to atmospheric Ar 36Ar/40Ari on the inverse isochron diagram (Fig X); effec tively avoiding the excess Ar apparent in other steps. These results suggest that SL-38 cooled below ~ 200-250C (approximate low temperature closure for K-feldspar) by abou t 57 Ma. This also indicates that the SLS in the westernmost transect was not reheated above ~200-250C during the Eocene. 40Ar/39Ar Thermochronology From Outside the Lower Plate Transect As noted above, four samples were collected for 40Ar/39Ar thermochronology from areas outside the transect across the northeastern Anaconda-Pintlar Range (Table 2).

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134 Two samples were collected from the LVC se quence within the ACC upper plate. DF02113, a crystal-lithic rhyo litic tuff, was collected from the Deerlodge Valle y, southeast of the eastern transect. Biotite from this sa mple yielded a weighted plateau age of 53.7 1.4 Ma (66% 39Ar gas, MSWD = 0.71). Note this bi otite age is slightly older than the biotite age of 51.4 0.4 Ma reported for a sample collected from the same outcrop (Isopolatov, 1997). A large, isolated granodi orite block was sampled (DF02-114) from within the LVC field, northeast of the DF02-113 sample localit y. This granodiorite is a medium grained biotite-hornblende bearing gr anodiorite similar to the SLS and other granodiorite intrusions in the ACC lower plate. Biotite from DF02-114 yielded a weighted plateau age of 76.3 1.1 Ma over 95% 39Ar gas release (MSWD = 0.57). Notably, this biotite cooling age is similar to that of WG04-114 (74.0 0.9 Ma), the SLS sample from the westernmost transect. The two remaining samples were collect ed from ACC lower plate exposures directly north of the eastern most transect in the Flint Creek Range. DF02-118a was collected from an outcrop of mylonitic biotite granite along Lost Creek, in the southernmost Flint Creek Range. Biotite from this sample gave a weighted plateau age of 38.8 1.62 (72% 39Ar gas, MSWD = 0.82). This bi otite cooling age is within analytical error of the biotite cooling age from DF02-120 (total fusion age = 38.5 0.7 Ma), the mylonitic biotite granite from Mill Creek located directly to the south. DF02119a was collected from two-mica granite of the Mt. Powell Bathol ith (granodiorite) on the eastern flank of the Flint Creek Range. Biotite and muscovite from this sample yielded slightly discordant pl ateaus. However, useful age information can be deduced from the results. The muscovite plateau of DF02-119a is characterized by four

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135 anomalous steps near the middle of the experime nt (Fig. 3). With these steps removed an error plateau age of 68.1 2.77 Ma was calcu lated from the remaining steps (55% 39Ar gas, MSWD = 4.33). The biotite plateau of DF02-119a is a slightly convex-up shaped, where the first and last few steps of the experi ments resulted in sli ghtly younger ages. By excluding these steps, a wei ghted plateau age of 68.8 1.40 Ma was calculated over 45% of the 39Ar gas release, corresponding to six co ntiguous steps from the middle of the experiment (Fig. 3). Note these mica cooli ng ages are slightly ol der than those reported for the Mt. Powell Batholith (~ 61.559.7 Ma) by Martin et al. (1989).

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136 CHAPTER 7 DISCUSSION Origin of the Lake of the Isle Shear Zone Field observations (Chapter 3) indicate that at leas t two phases of high grade metamorphism have affected the Lake of the Is le shear zone (LISZ): (1) an early higher pressure event shown by fresh and relict kya nite in the Lower Belt Greyson Formationequivalent (Kalakay et al., 2003) and pelitic upper Ravalli Group-equivalent metasedimentary strata and (2) a later lower pressure and high temperature upperamphibolite facies event that created sillim anite-bearing assemblages which overprint the earlier higher pressure assemblage in some ar eas. Localized anatexis (in pelitic strata) and significant ductile attenuation of Belt and Middle Cambrian-equivalent apparently accompanied the later upper-amphibolite facies metamorphic event (Chapter 3). Here, new age constraints obtained from UPb geochronology are discussed for upperamphibolite facies metamorphism and the a ssociated localized anatexis and ductile deformation in the LISZ. In addition, pr essure-temperature constraints obtained by conventional thermobarometry for the uppe r-amphibolite facies metamorphism are discussed in context of the PT history of the LISZ. Together these constraints, combined with field observation, are used to provide a kinematic history for the LISZ. The age and pressure-temperature constraints obtained fo r upper-amphibolite facies metamorphism in the LISZ are also used in a later section as a proxy for estimating the maximum amount of exhumation in the ACC facilitated by extension during the Eocene (see below).

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137 Age Constraints The most direct age constraint for uppe r-amphibolite facies metamorphism, local anatexis, and ductile deform ation of meta-Belt and meta -Cambrian metasedimentary strata in the LISZ would be a U-Pb zircon crystallization age from granitic leucosome sample WG05-02. However, all the zircons anal yzed from this sample are interpreted to be Proterozoic and Archean zirc ons inherited from the protolith (meta-Greyson) that have undergone variable lead loss (Chapter 4). De spite the lack of relevant age information obtained from the leucosome zircons the U-Pb zircon ages obtained from the undeformed Storm Lake Stock (SLS) granodiorite and the de formed quartz diorite sill, combined with some important field observations, provide substantial age constraints for upper amphibolite facies metamorphism and the relate d events in the LISZ. Field relationships are summarized below, followed by a discussi on of the U-Pb ages obtained from these intrusions and implications for the age of the upper-amphibolite facies metamorphism, local anatexis, and ductile deformation in the LISZ. In the central and eastern part s of the LISZ, the deformed quartz diorite sill exhibits a solid-state foliation concordant with th e transposed metamorphic foliation in the adjacent metasedimentary strata (Chapter 3, also see Appendix B and D). This relationship indicates the emplacement of the qua rtz diorite sill either (1) predated high grade metamorphism and ductile deformation in the LISZ or (2) was synchronous with these events. In the western LISZ, the unde formed SLS granodiorite clearly cross-cuts deformed upper-amphibolite-grade meta-Be lt and meta-Cambrian metasedimentary strata. In addition, the undefo rmed quartz diorite of the SL S (not dated in this study) cross-cuts these deformed metasedimentary strata; the contact between these two intrusions shows the quartz di orite phase to be earliest (A ppendix B and F). Granitic

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138 leucosome is commonly observed in the meta -Greyson pelitic para gneiss and often found in boudin neck pressure shadows between th e boudins. The longer axes of the boudins are oriented parallel to th e metamorphic foliation in the metasedimentary country rock and deformed quartz diorite sill indicati ng local anatexis in the meta-Greyson was synchronous with ductile deformation in the LISZ (Chapter 3). Together, these field observations indicate the followi ng relative order of events in the LISZ: (1) emplacement of the quartz diorite sill before or sy nchronous with (2) upper-amphibolite facies metamorphism, local anatexis in the meta -Greyson pelitic para gneiss, and ductile deformation of metasedimentary strata fo llowed by (3) emplacement of the SLS quartz diorite and then (4) the emplacemen t of the SLS granodiorite. 206Pb / 238U weighted mean zircon ages obtaine d from the deformed quartz diorite sill and SLS granodiorite phase are 75.0 0.8 and Ma 74.6 0.8, respectively (2 errors, Chapter 4). The concordance of these crystal lization ages indicate th e quartz diorite sill and SLS granodiorite were emplaced into the LISZ at a very similar time. Because of the cross-cutting field relationshi ps described above, the SLS quartz diorite in the western LISZ must have been emplaced at a similar time, presumably during the short time interval between emplacement of the quartz diorite sill a nd the SLS granodiorite. Because the 206Pb / 238U weighted mean zircon ages of the deformed quartz diorite sill and undeformed SLS granodiorite are within analytical error it is not likely that the emplacement of the quartz diorite sill signi ficantly predated upper-amphibolite facies metamorphism, local anatexis, and ductile deformation in the LISZ. Instead, emplacement of the quartz diorite sill was pr obably synchronous with some stage of these events. The outcrop pattern of migmatites within pelitic Greyson Formation-equivalent

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139 paragneiss adjacent to the quartz diorite sill in the central LISZ supports this interpretation (Appendix F). The migmatites only occur directly adjacent to the sill ( 0.5 km) which indicates the emplacement of the sill provided a thermal pulse causing localized anatexis and more significant duc tile attenuation of adjacent meta-Belt and meta-Cambrian metasedimentary strata (see below). Because the quartz diorite sill exhibits a solid-state foliation that is conc ordant to the metamorphic foliation in the adjacent metasedimentary strata, ductile deformation must have continued for at least some time after emplacement of the sill. Emplacement of the undeformed SLS quart z diorite, followed by the emplacement of the SLS granodiorite, must have occurred immediately after ductile deformation had ceased in the LISZ because these two intrusi ons cross-cut the deformed metasedimentary strata. The thermal affects of these intrusio ns on metasedimentary strata in the western LISZ has not yet been determined. However, a biotite cooling age of 74.0 0.9 Ma from near the center of the SLS granodiorite plut on (WG04-114) show this intrusion, and at least the western LISZ, rapidly cooled to temperature less than ~350C by about 74 Ma (Chapter 6). Upper-amphibolite facies metamorphism and synchronous ductile deformation had ceased before ~74 Ma, at least in the western LISZ. In summary, the U-Pb geochronology, 40Ar/39Ar thermochronology, and field relationships summarized here show upper-a mphibolite facies metamorphism, local anatexis in the meta-Greyson pelitic paragne iss, and ductile atte nuation in the LISZ occurred from ~75-74 Ma. More constraint s may be obtained for the duration of these events by: (1) Use catholuminescence imaging (CL) to identify magmatic rims on the WG05-02 leucosome zircons from the migmatitic pelitic Greyson Formation-equivalent

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140 paragneiss. If thin magmatic rims are pres ent, then employ a higher resolution (<30 m) beam (e.g., SHRIMP) to obtain U-Pb isotopic da ta and ages from the rims. (2) Separate and date metamorphic zircon and/or titanite from the attenuated upper-amphibolite facies metasedimentary strata in the LISZ usi ng U-Pb geochronology (e.g., Dziggel, et al., 2005). Pressure-temperature History Relict and fresh kyanite-bearing assemblages in some parts of the LISZ indicate earlier higher pressure metamorphism prio r to the lower pressu re, high temperature metamorphism at ~75-74 Ma (Chapter 3). The earlier higher pressure metamorphism in the LISZ may have followed a Barrovian-sty le clockwise PT path related to tectonic loading during emplacement of the Sapphire thru st slab in the Late Cretaceous. Peak pressures associated with the early metamo rphism were probably no more than ~7-8 kbar, based on a reconstructed thickness of ~ 20-25 km for the Sapphire thrust slab in the Late Cretaceous (Sears and Hendrix, 2004, see cr oss sections in their Fig. 2 and 4) and that the LISZ was probably near the base of the Sapphire thrust sl ab (Wallace et al., 1992). Although this may be a viable hypothesis, the age and extent of earlier higher pressure metamorphism in the LISZ is curre ntly unknown. Therefor e, only the pressuretemperature history of the LISZ during the later lower pressure and high temperature metamorphism is further discussed. The thermobarometric data obtained from the LISZ in this study (Chapter 5), combined with field relationships an d metamorphic petrology, indicate the metasedimentary strata in the LISZ underw ent prograde isobaric heating during the low pressure, high temperature metamorphic event at ~75-74 Ma. Peak metamorphic conditions during this metamorphism were at the upper-amphibolite facies at ~3.2-5.3

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141 kbar and ~750-825C (Chapter 5). The marked increase in metamorphic grade within the Meta-Greyson pelitic metasedimentary strata towards the center of the LISZ, and the contact with the deformed quartz diorite si ll, indicates the thermal peak during this metamorphic event is directly related to a heat from emplacement of the sill at ~75-74 Ma. Therefore, the mineral/textural zone s and zone boundaries mapped in the Greyson Formation-equivalent pelitic strata were co rrelated with specific divariant fields and invariant metamorphic reactions in PT space (Table 3-1, Fig. 3-16). Combining these empirical constraints with the best pressu re-temperature estimate obtained from the thermobarometry of this study (Chapter 4), a PT path can be constructed for the LISZ during the later low pressure, high temperature metamorphism. Figure 7-1 displays a phase diagram for the KFMASH chemical system representative of pe litic rocks along with the peak pr essure-temperature estimate for sample ME-231, select labeled divariant fiel ds and univariant reactions, and the proposed isobaric PT path (Path A) for the LISZ during the low pressure, high temperature metamorphism. At distances greater than ~1 km from the deformed quartz diorite sill contact, within mineral/texture zone 1, the Meta-Greyson pelitic schi st contains; primary muscovite, abundant coarse-grained biotite, garnet and K-feldspar porphyroblasts, and fine-grained sillimanite fibrolite. This asse mblage corresponds to divariant field 1 (D1) in Figure 7-1, immediately up temperature of the reaction staurolite + chlorite aluminosilicate + biotite (7-1) which is responsible for the appearance of aluminosilicate wi th heating during prograde metamorphism (Spear, 1993); the type of aluminosilicate polymorph produced across reaction 7-1 is solely dependent s on the pressure (Fig. 7-1).

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142 Figure 7-1. Phase diagram showing the Late Cretaceous PT history of the Lake of the Isle shear zone (LISZ). The numbere d univariant reactions and divariant fields (D1 and D2) are discussed in the text. Path A represent the isobaric PT path proposed for the LISZ during lower pressure high temperature metamorphism associated with the empl acement of the quartz diorite sill. The dashed PT path corresponds to the poor ly constrained earlier higher pressure metamorphic event that may have been associated with tectonic loading during emplacement of the Sapphire thrust plate (see text). Closer to the deformed quartzdiorite sill, at distances between ~1-0.5 km, the MetaGreyson pelitic schist lie in mineral/textu ral zone 2 and contains; relatively coarsergrained sillimanite fibrolite mats, K-feldspar garnet, abundant biot ite and lack primary muscovite suggesting these rocks passed thr ough the muscovite-out dehydration reaction muscovite aluminosilicate + K-feldspar + H2Ovapor (7-2) during prograde heating and fall in divariant fi eld 2 (D2) in Figure 71. Still closer to the deformed quartz diorite sill contact, at distances of 0.5 km, the Meta-Greyson pelitic

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143 schist grades to highly deformed paragnei ss and corresponds to mineral/texture zone 1 with an upper-amphibolite facies assemblage comprised of; thick relatively coarsergrained sillimanite fibrolite mats, less a bundant and altered biotite, large sigmoidalshaped K-feldspar and cordierite porphyrobl asts, large garnet porphyroblasts, common granite leucosome, and lack primary muscovite. Note that the grani tic leucosome can bee seen cross-cutting garnet porphyroblasts and with in cordierite pressure shadows in thin section (Fig. 3-15). Together, these featur es indicate that the meta-Greyson pelitic paragneiss directly adjacent to th e deformed quartz diorite sill ( 0.5 km), within mineral/texture zone 1, passed through the reactions biotite + aluminosilicate garnet + cordierite + H2Oliquid (7-3) K-feldspar + garnet + cordierite + biotite + H2Ovapor liquid melt (7-4) during prograde metamorphism. Note reacti on 7-4, the pelite minimal melting reaction, requires the presence of a vapor phase to pr oduce a melt phase (Spear et al., 1999). Some vapor phase would have been necessary to cause anatexis of the meta-Greyson pelitic paragneiss near the quartz di orite sill at near peak temperature conditions during metamorphism. One possibl e vapor source could be H2O produced by the dehydration of biotite as reaction 7-3 was cr ossed immediately before cros sing reaction 7-4 or 7-2, see Fig. 7-1). Passage through reaction 7-3 and 7-4 with production of granitic leucosome during anatexis is consistent with the pr essure-temperature estimate of ~3.2-5.3 kbar, ~750-825C obtained for sample ME-231, the meta-Greyson migmatitic paragneiss sample located ~0.4 km from the contact with the deformed quart diorite sill. Following the thermal peak in upper-amphi bolite facies metamorphism, the MetaGreyson pelitic schist and paragneiss most likely underwen t isobaric cooling following

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144 the prograde path in reverse. Upon cooli ng, and passing through reaction 7-4 in reverse, all leucosome melt present would have crystallized and any H2O release associated with crystallization of the leucosome would have ceased on the low temperature side of reaction 7-4; at these temperatures most free H2O would have likely been expelled to shallower depths (Spear et al., 1999). Th e absence of a signifi cant quantity of vapor during further cooling would have prohibite d the prograde reactions 7-2 and 7-3 from operating in reverse. Therefore, no retrogr ade muscovite would be produced as reaction 7-2 was crossed during cooling. This is consis tent with the absence of muscovite in the Greyson Formation-equivalent pelitic migmatitic paragneiss located directly adjacent to the deformed quartz diorite sill. The presence of fine-grained sillimanite fibrolite in large cordierite porphyroblast also supports the reve rsal of reaction 7-3 during isobaric cooling (Table 3-1, Fig. 7-1). Kinematic Interpretation The new U-Pb geochronology, thermobarometr y, and field observations show that low pressure, high temperature metamorphism in the LISZ occurred synchronous with emplacement the quartz diorite sill at ~7574 Ma. The low pressure, high temperature apparently posted higher pressure kyanite -zone metamorphism event in the LISZ; the origin of the earlier metamorphism is not ye t constrained. Field obs ervations (Chapter 3) show that ductile attenuation of metasedime ntary strata in the LISZ accompanied the upper-amphibolite facies metamorphism and lo calized anatexis. The most convincing line of evidence for the coincidence of these events is the common occurrence of granitic leucosome within pressure shadows betw een flattened mesoscopic boudins in the intensely deformed meta-Greyson pelitic pa ragneiss adjacent to the deformed quartz diorite sill (Fig. 3-9). Emplacement of the gr anodiorite and quartz diorite phases of the

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145 SLS must have posted dated the ductile deformation because these intrusions are undeformed and obliquely cross-cut the ductile metamorphic fabrics in the metasedimentary strata. Strain localization and ductile attenuation of metasedimentary strata in the LISZ Low pressure, high temperature metamorphism associated with emplacement of the quartz diorite sill into the LISZ at ~75-74 Ma resulted in a thermal anomaly that surrounded the deformed quartz diorite sill. This is shown by the marked changes in metamorphic assemblages and textures in the metamorphosed Greyson Formation perpendicular to the contact with the sill (see Table 3-1 and Fig. 3-16). The thermal anomaly surrounding the sill apparently locali zed ductile attenuation of Belt and Middle Cambrian-equivalent metasedime ntary strata within the LISZ, as the degree of ductile strain (attenuation) increases towards the sill contact in the center of the LISZ. The thermally induced ductile strain gradient is indicated by the marked change in the geometry of mesoscopic folds and boudins towa rds the sill; folds are increasing isoclinal and recumbent and boudins are progressively flattened towards the sill contact and the center of the LISZ (Chapter 3). The mark ed change in the geometry of mesoscopic structures towards the sill and center of the LISZ (and attenuation) is the result of major fabric transposition under ductile flow wh ich rotated, stretched, and flattened the mesoscopic structures into an orientation parallel to the me tamorphic foliation or gneissic banding in the host metasedimentary strata a nd the contact with th e quartz diorite sill (e.g., Davis and Reynolds, p. 453). Therefore, ductile attenuation of Belt and Middle Cambrian-equivalent metasedimentary strata in the LISZ was fac ilitated by a reduction rheology (rock strength) caused to the thermal pulse related to empl acement of the quartz diorite sill at ~75-74 Ma.

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146 Strain classification The transposition of structures and fabr ics and attenuation of the Belt and Middle Cambrian metasedimentary strata in the LI SZ was accommodated predominately by pure shear and plane (flattening) coaxial stra in (e.g., Passchier et al., 1993, p. 16). The predominance of plane coaxial strain in the LISZ is indicated by the widespread occurrence of symmetric porphoroblasts shear indicators. However, documentation of asymmetric porphyroblasts showing a left late ral sense of shear indicate in some areas ductile deformation was accommodated, at least in part, by localized left lateral simple shear and noncoxial strain (C hapter 3). These areas pr obably represent small-scale sinistral shear zones but the significant of thes e structures has not yet been determined. It is possible these smaller shear zones accomm odated differential strain in the LISZ. Structural Interpretation In the central part of the LISZ, near th e Lake of the Isle, Lower Belt Meta-Greyson schist and paragneiss are j uxtaposed with the Middle Cambrian-equivalent strata separated only by the deformed quartz diorite sill. To the west, towards Storm Lake, progressively younger Belt-equivalent metasedime ntary strata in the south are juxtaposed with the Middle Cambrian-equivalent metasedime ntary strata to the north. However, in the western LISZ, the undeformed SLS (Sto rm Lake Stock) granodiorite and quartz diorite intrude the shear zone and separate these juxtaposed metase dimentary strata by a significant distance (see geologi c map, Appendix D). The fact the apparent offset between Belt and Middle Cambrian metase dimentary across the LISZ progressively decreases to the west permits the shear zone to be interpreted as a fault which cuts up section towards the west. In the central LISZ, the juxtaposition of the Lower Belt MetaGreyson with Middle Cambrian-equivalent strata indicates an enormous amount of offset.

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147 In the west the entire metamorphosed Ravalli Group to Middle Cambrian section is apparently repeated but separa ted by the voluminous SLS. The structure is complex in the eastern LISZ Here, the LISZ consists of the Mill Creek nappe, a km-scale east-verging recumben t fold originally mapped by Emmons and Calkins (1913) and then in more detail by Heise (1983). The Mill Creek nappe folds the metamorphosed and attenuated Helena Forma tion through meta-Cambrian section and is cored by a thrust fault which places the Helena Formation over the metamorphosed Middle Cambrian Silver Hill Formation (Appendi x F). Presently, it cannot be shown the LISZ is a thrust or normal fault structure. However, the Late Cretaceous age of the LISZ and the fact that several plut ons have intruded into the structure (similar to many thrusts faults elsewhere in the Sapphire thrust plate) combined with thrust fault structure in the Mill Creek drainage suggest the entire LISZ is a thrust fault. Eocene Exhumation and Cooling History of the Anaconda Metamorphic Core Complex defined by 40Ar/39Ar Thermochronology The 40Ar /39Ar thermochronological data set obtai ned from rock samples collected across the ACC lower plate in the current study area define a lateral cooling age gradient, where cooling ages young progressively to th e ESE across the lower plate. Individual mineral cooling ages and transect sample lo calities are shown in Figure 7-2 plotted on a simplified geologic sketch map. In addition, mica cooling ages (muscovite and biotite) have been contoured in Figure 7-3 to clearly illust rate the lateral cooling age gradient across the exposed ACC lower plate rocks. The presence of the lateral cooling age gradient and the fact that cooling ages young in same direction as tectonic transport (detachment) of the upper plate (ESE, 110-100 as defined by stretching lineations and kinematic indicators in the greenschist myl onites, Kalakay et al ., 2003) indicates the

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148 dominate mechanism for vertical exhumati on and cooling of the ACC lower plate was tectonic unroofing, not erosion (Foster et al., 1990; McGrew and Snee, 1994; Foster and John, 1999). Thermochronological data obtained by 40Ar/39Ar thermochronology in this study are discussed below in the context of the tectonic exhumation and cooling of the ACC lower plate. In particular, the new thermo chronological data combined with previous thermochronology and geochronology provide cons traints on: (1) the thermal history of the ACC lower plate, (2) the onset and duration of extension in the ACC, (3) the slip rate and geometry of the bounding normal detachment fault, and (4) the magnitude of offset facilitated by the bounding detachment fau lt since the onset of extension. Lower Plate Cooling History The cooling history of the ACC lower pl ate exposed in the current study area is now defined from the late Cretaceous to the Eocene by new 40Ar /39Ar thermochronological data, U-Pb geochronology, and thermobarometric data obtained in this study combined with previous thermochronology a nd U-Pb geochronology. All relevant data has been compiled and su mmarized in Figure 7-4 which displays a temperature (y-axis) vs. time (x-axis) coo ling diagram for the ACC lower plate. The temperature-time cooling diagram contains fiel ds representing the approximate range in temperature and age for magmatic and meta morphic events recorded by the ACC lower plate since late Cretaceous time. In additi on, cooling curves have been constructed for major intrusions emplaced into the ACC lower plate and for different regions of the ACC lower plate country rock. During the late Cretaceous much of the A CC lower plate was at temperatures in

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149 Figure 7-2. Geologic sketch map showing th e mineral cooling ages obtained by 40Ar39Ar thermochronology from samples collected along the ACC lower plate transect in this study. Mine ral cooling ages are reported with two sigma error.

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150 Figure 7-3. Mica cooling age contou r map constructed from biotite and muscovite cooling ages obtained 40Ar39Ar thermochronology from samples collected along the ACC lowe r plate transect in this study. Mica co oling ages young to the east across the footwall, consistent with top-to-the-east-southeast Eo cene unroofing of the lower pl ate (see text for details).

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151 excess of ~600-700C. Thermobarometric da ta obtained from migmatitic meta-Greyson pelitic gneiss in the LISZ i ndicates peak metamorphic te mperatures reached ~750-850C during upper-amphibolite facies metamorphism which occurred no later than ~74-76 Ma (see Chapter 5). Previous pressure-temperatu re estimates from other regions of the ACC lower plate (e.g., the Flint Creek and southern Anaconda-Pintlar Ranges) suggest similar metamorphic temperatures at this time (s ee Chapter 5). Emplacement of the several voluminous batholiths, stocks, and plutons at ~80-65 Ma (e.g., the Flint Creek plutons, Storm Lake Stock, and Boulder and Pioneer Ba tholiths, see Fig. 2-2 and 2-4) maintained high temperatures in the ACC lower plate during the late Cretaceous (Fig. 7-4). Figure 7-4. Temperature-time cooling diagra m showing the cooling history of the ACC lower plate. The diagram was constr ucted from the new and previous thermochronology, geochronology, and thermobarometry from the ACC lower plate.

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152 The ACC lower plate underwent rapid cooling immediately following upperamphibolite facies metamorphism in the LISZ and emplacement of voluminous intrusions during in the late Cretaceous. However, th e lateral gradient in mica cooling ages now documented across the exhumed lower plate w ithin the current study area (Fig. 7-2 2 and 3) indicates the eastern and western parts of the ACC lower plate underwent substantially different cooling histories upon tectonic exhum ation. The difference in cooling history, at least from the middle Eocene on, is due to tectonic exhumation of the lower plate along the east-dipping detachment (i.e., the easte rn ACC lower was down-dip and deeper [and warmer] than the western ACC lower during ex tension). In addition, the eastern (deeper) part of the ACC lower plate was intruded by gr anites and granodiorite plutons in the early to middle Eocene time causing its cooling hist ory to substantially differ from the western (shallower) part of the ACC lower plate. The cooling history of th e western lower plate is now defined by geochronology and thermochronology from the Storm Lake St ock (cooling curve 1, Fig. 7-4). Following emplacement of the quartz diorite and granodi orite phases of the SLS at ~75-74 Ma (see Chapter 4) the western lower plate cooled ve ry rapidly to less than ~350C by ~74 Ma based on a well defined biotite cooling age from the SLS gra nodiorite (see Chapter 6). The ~350C closure temperature fo r this biotite was calculated for rapidly cooled sample. Subsequent to ~74 Ma, however, cooling of the SLS and western ACC lower plate was slower. A low-temperature error plateau obtai ned from laser step-heating of K-feldspar from the SLS granodiorite gave a 40Ar/39Ar cooling age of ~57 Ma indicating the western lower plate most likely remained above ~250-200C until ~57 Ma. After ~57 Ma, the low temperature history of the western part of the ACC lower plate cannot be further

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153 constrained with low-temperature thermochr onological data such as apatite fission track or U-Th/He cooling ages. The eastern part of the ACC lower plate slower following upper-amphibolite facies metamorphism and emplacement of late Cretaceous intrusions compared to the western part of the lower plate (coo ling curve 2, Fig. 7-4). Presumably, the eastern ACC lower plate initially cooled slower following the la te Cretaceous thermal maximum because this part of the lower plate was at deeper and th erefore at higher ambient temperatures during this time. At ~53-47 Ma two-mica and biotit e granites and granodior ite plutons intruded the eastern (deeper) part of ACC lower plate based on 206Pb/238U zircon crystallization ages (Foster et al., 2006a). These intrusions most likely cooled rapi dly at first because the elevated geothermal gradient in the A CC lower plate during the late Cretaceous (~4050C/km) had since decayed, probably to ~ 25-30 C/km (cooling curve 3, Fig 7-4). However, biotite and muscovite from these in trusions gave cooling ages of ~40-39 Ma indicating the intrusions, and the eastern lower plate countr y rock, remained above ~350400C before this time. Muscovite and biotit e cooling ages from these intrusions are within error indicating the onset of rapid cooling ( 125C/Ma) at ~40-39 Ma most likely due to a change in the detachment geometry in the upper crust (see below). The cooling history of the eastern lower plate within the current study area cannot be further constrained at this time without lower thermochronological data (e.g., apatite fission track or U-Th/He cooling ages). Fo ster and Raza (2002) report apatite fission track cooling ages from the eastern part of the Chief Joseph Batholith in the southern Anaconda-Pintlar Range. These cooling ages range from ~40-30 (younging approximately west to east) and record th e time at which the Chief Joseph Batholith

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154 cooled below ~120-90C within the shallow crust (Reiners et al., 2005). These low temperature thermochronological data approxi mate the low temperature cooling history of the eastern ACC lower plate, at least in the southern part of the exposed lower plate (dashed part of cooling curve 2, fig 7-4). Constraints on the Timing of the Onset of Extension The timing of the onset of extension in fault-bound metamorphic core complexes can be constrained using lower plate 40Ar/39Ar thermochronology in two ways. One strategy is to document the timing of acceler ated cooling of the lower plate, which presumably commences at the onset of rapid ex tension; lower plate ro cks cool relatively slow prior to extens ion and then quickly during rapid extension as they are uplifted towards the surface (e.g., Foster and John, 1999; other references?). This method, referred to as the accelerated cooling met hod here, requires temperature information from individual lower plate rock samples pr ior to, during and afte r the onset of rapid extension therefore spanning a wide range of temperatur e and time (e.g., U-Pb zircon crystallization ages, hornblende, mica, K-feld spar and apatite fissi on-track cooling ages, ranging ~750-100C, their respective closure temperatures). Following this approach, mineral cooling ages, sometimes from multiple lower plate samples, are typically plotted on a temperature (y-axis) vs. time (x-axis) diag ram and cooling curves are fit to the data. A sharp increase (or break) in the slope of the cooling curve for the samples marks the time at which exhumation began, causing rapid cooling (Foster and John, 1999). Another way to constrain the timing of the onset of extension in metamorphic core complexes using lower plate 40Ar/39Ar thermochronology is to document the position of preserved or quenched palaeoisotherms with in the exhumed lower plate (Foster et al., 1993; Foster and John, 1999). In particular the position of the palaeoisotherm

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155 representing the base of the partial retention zone (PRZ) fo r a given temperature sensitive mineral system can be used to estimate the ag e of the onset of extension (Foster and John, 1999). This approach is referre d to as the PRZ method here. The PRZ refers to a zone (or range of depths) within a vertical column of crus t where heating, due to the geothermal gradient, causes the daughter pr oduct for a given radiogenic decay scheme to be only partly retained by its host mine ral (Stockli, 2005). This occurs because temperatures within the PRZ progressively a pproach the closure te mperature of a given mineral with increasing depth (e.g., the Ar clos ure temperatures of K-bearing minerals). Above the PRZ, in the case of the K-Ar syst em, K-bearing minerals record cooling ages unaffected by heating from the geothermal gradie nt at a given time, presumably related to earlier cooling events. Within the range of depths defining the PRZ, cooling ages from K-bearing minerals are considered mixed ages because some amount of 40Ar is not retained during radiogenic decay of 40K due to heating with in creasing depth (Foster and John, 1999). However below the PRZ, K-bearing minerals are effectively zero aged because at these depths the crust is too warm to retain any 40Ar produced by radiogenic decay. Only when these deeper rocks are rapidly uplifted through the PRZ will Kbearing minerals within the rocks acquire coo ling ages. Thus, at th e onset of extension, for example in metamorphic core complexes, rocks at the base of the PRZ are uplifted along detachment faults through the PRZ and record cooling ages that correspond the age of the onset extension. Rocks at deeper crustal levels (i.e., significantly below the base of the PRZ) will record cooling ages progre ssively younger than the onset of extension because these deeper rocks are uplifted through the PRZ at progressively later times (Foster and John, 1999; Stockli, 2005).

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156 To estimate the timing of the onset of ex tension in metamorphic core complexes using the PRZ method as described here the fo llowing steps must be ta ken. (1) A suite of rock samples should be collected along a lower plate transect in the direction of slip on the bounding detachment fault system(s). Therefore, the sample s collected along the transect will young in the slip direction and represent once progressively deeper rocks (i.e., increasing paleodepths), given that the lower plate ro cks were exhumed by tectonic unroofing (Foster and John, 1999; St ockli, 2005). (2) A cooling age (y-axis) vs. distance in slip direction (x-axis) diagram is then constructed from cooling ages obtained from thermochronology; the distance in slip directi on corresponds to the relative positions of the rock samples along a chosen lower plate tr ansect line. A curve is then fit to the cooling age data for a given mineral system (3) The base of the PRZ for a given temperature sensitive mineral system can then be identified by locating an inflection point defined by a dramatic change in the slop e of the cooling age curve in the direction of slip. The inflection point is usually ma rked by a change from sub-vertical to subhorizontal slope of the curve in the slip dire ction. (4) A horizontal line is then passed through the inflection point and will intersec t the y-axis at the age of the onset of extension. A cooling age from a rock sample th at falls along this line is equal to the age of the onset of extension because it was at the base of the PR Z. Rock samples that fall below the line record cooling ages after the onset of extension because they were below the base of the PRZ prior to extension. Mo st rock samples above this line represent lower plate rocks that were within the PRZ prior to the onset of extension and record mixed cooling ages. However, the oldest ro ck samples near the top of the diagram can represent cooler, shallower rocks that were above the PRZ before extension began. Note

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157 also a vertical line passed through the infl ection point in the age vs. distance in slip direction diagram will intersect the x-axis at the distance along the lower plate transect line representing the base of the ex humed PRZ (Foster and John, 1999). In this study, the PRZ method was utilized to constrain the timing of the onset extension and exhumation of the ACC lowe r plate. The accelerated cooling method could not employed because the thermochr onological data obtained from individual lower plate samples do not span the temperature-time range needed to employ this method (i.e., from pre-to-syn-to-post-extensi onal temperatures). Below, the PRZ method is applied to the mica cooling ages obtained along the ACC lower plate transect in the current study area. Figure 7-5a displays an age vs. distance in slip direction diagram constructed from muscovite and biotite cooling ages obtained from lower plate rock samples using 40Ar/39Ar thermochronology. In addition, Figu re 7-5b shows a simplified geologic stretch map of the exhumed ACC lower plate in the current study area with rock sample localities and the chosen transe ct line (oriented at 105, the average slip direction on the detachment). Rock samples adjacent to th e transect line were projected orthogonally back to the line defining their distances in slip direction as show n along the x-axis in Figure 7-5a. Horizontal (dista nce) error bars for each rock sample are 1 km. It is common practice to apply distan ce in slip direction error on these types of plots to account for errors associated with sample proj ections and variability in elevation (i.e., paleodepth) between transect rock samples (F oster et al., 1993; Brichau et al., 2005). Vertical error bars represent errors in indivi dual mica cooling ages. The distance in slip direction and age error bars in Figure 5a represent 2 errors (95% confiden ce interval).

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158 Both muscovite and biotite cooling ages progressively young across the ACC lower plate to the ESE, in the direc tion of slip on the detachment fault system (Fig. 7-1,7-2, and 7-5a). This trend in cooling ages is consis tent with the interpretation that rock samples collected in the direction of slip across exhumed lower plates represent increasing paleodepths (e.g., Fitzgerald, et al., 1991; Foster et al., 1993; Lee, 1993, Foster and John, 1999). Therefore, these thermochronological data can be used to constrain the position of the exhumed PRZ for micas and the age of th e onset of extension in the ACC. At a distance of ~4-5 km along the lower plate tr ansect mica cooling ages drop from Late Cretaceous ( 74 Ma) to middle Eocene cooling ages ( 53 Ma). Further to the ESE, in the direction of slip, mica cooling age more gradually decrease to ~39 Ma. The marked change in slope of the mica cooling age curv e defines the position of the palaeoisotherm that corresponds to the base of the exhumed PRZ for micas (F ig. 7-2 5a, b). The top of mica PRZ is constrained to be just ESE of (or slighter deeper in paleodepth than) WG04114, a rock sample collected from the SLS gr anodiorite (Fig 7-5a, b). This sample yielded a biotite cooling ag e of 74.0 0.9 Ma and a zircon 206Pb/238U weighted mean age of 75.6 1.1 (Chapter 4). The concordance of these ages indicates the SLS granodiorite underwent rapidly post-magmatic cooling after emplacement into cooler, shallow crust above the mica PRZ. Therefore, samples WG04-052 and WG04-112 lie within the exhumed mica PRZ (Fig. 7-5a, b). A horizon tal line passed through the mica cooling age inflection point intersects the y-axis at ~53 Ma the age of the onset of extension in the ACC and initial exhumation of the lower plate (Fig. 7-5a).

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159 A B A B Figure 7-5. Age vs. distance in slip direc tion diagram and lower plate transect sketch map. A) The age vs. slip rate dire ction diagram was construct by projecting samples back to a defined transect line which has been oriented parallel to the average slip direction on the detachment (105). PRZ = partial Ar retention zone for micas. B) The lower plate transect sketch map shows the sample projections used to construc t the age vs. slip rate direction diagram. Samples projected back to the transect line wi th dashed lines were not used in subsequent slip rate calculations. Constraints on the Duration of Extension The duration of extension and tectonic exhumation of the ACC are now largely constrained by lower plate mica cool ing ages obtained here and by 206Pb/238U zircon

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160 crystallization ages from the granitoid mylonites sampled from the greenschist facies mylonitic shear zone in th e easternmost study area (206Pb/238U zircon ages from Foster et al., 2006a). These data indicate that extens ion, causing exhumation and cooling of the ACC lower plate, began at ~53 Ma within th e mylonitic shear zone, presumably at midcrustal depths. Extension in the ACC, facilitated by ductile flow in the mylonitic shear zone and by brittle detachment at shallower crustal levels, continued until ~40-39 Ma based the cooling ages of mica from the gr eenschist facies mylonites. These mica cooling ages record the time when the easte rn part of the shear zone was exhumed upwards through the brittle-duc tile transition (at ~300-350C) therefore causing ductile deformation to cease. ONeill et al. (2004) report a cooling age of 47.2 0.3 Ma (given with 2 error here) from a single muscovite porphyroclas t extracted from a sample of mylonitic micaeous quartzite in the Sullivan Creek drainage within the southern part of the current study area. In this study, a similar cooli ng age of 47.3 1.1 Ma was obtained from a fraction of muscovite porphyrocla sts extracted from a sample collected at the ONeill et al. (2004) sample site. These cooling ages most likely record recrystallization (or neocrystallization) of the preexisting muscovite porphyroclast s within the quartzite at or just below the closure temperature of muscovite in the latest stages of ductile deformation under greenschist facies conditi ons (Dunlap, 1997; Foster, et al ., in press). Thus, ductile deformation apparently ceased much earlier in this part of the mylonitic shear zone. The differences in age of ductile deformation in thes e two part of the shear zone is consistent with ESE directed tectonic unroofing of the AC C lower plate. Because this part of the mylonitic shear zone lies west (and was pr eviously up-dip on the detachment) of the

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161 shear zone exposed in the Mill and Clear Cr eek drainages it would have been exhumed through the brittle-ductile transition at an ea rlier time during extension. Apatite fission track cooling ages from the southern An aconda-Pintlar Range reported by Foster and Raza (2002) indicate that extension, exhuma tion, and cooling of the ACC lower plate continued until ~30-25 Ma, at least in the south. Constraints on the Detachment Slip Rate Thermochronological data from rock samp les collected along lower plate transects in metamorphic core complexes parallel to th e direction of tectonic unroofing have also been successfully used to estimate previous slip rates on bounding detachment faults (e.g., Foster et al., 1993; J ohn and Foster, 1993; Foster a nd John, 1999; Wells et al., 2000; Stockli, 2005). The mica cooling ages obt ained from the ACC lower plate transect in this study may be used to estimate the sl ip rate on the bounding de tachment fault. In order to make realistic and m eaningful slip rate estimates from these thermochronological data some important assumptions must be ma de. (1) Vertical exhumation and cooling of the ACC lower plate during the time interval ove r which the slip rate is estimated must have been due to tectonic unroofing (i.e., up lift by displacement al ong the detachment), not erosion. (2) Isotherms of the syn-ex tensional geothermal gradient must have remained approximately horizontal and stable over the time interval used to estimate a slip rate for the detachment (Ketchum, 1996; Foster and John, 1999; Stockli, 2005). In the case of the ACC lower plate, the first assumption is most likely valid. As noted above, the presence of a lateral th ermal gradient across the ACC lower plate defined by the ESE younging trend in mi ca cooling ages precludes any major contributions of erosion to the exhumation of the ACC lower plate; th is holds true for at least the time interval in which the previ ously zero aged lower plate micas were

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162 uplifted and cooled below their Ar closure temperatures (from ~51-39 Ma). The second assumption, a stable syn-extensional geotherm al gradient, is also probably valid for the ACC but requires more discussion here. One potential cause for an unstable geothermal thermal gradient during extension in metamorphi c core complexes is heat advection. For example, when deeper and warmer rocks ar e rapidly uplifted along a detachment fault they are placed directly adjacent to relatively shallow and cooler rocks. As a result, heat from the previously deeper rocks will advect into the surrounding shallower, cooler rocks (i.e., cooling of the deep rocks does not k eep the pace with tect onic uplift, House and Hodges, 1994; Scott et al., 1998). Consequently a lateral thermal gradient is created across the detachment fault possibly resulting in an unstable geothermal gradient during extension characterized by non-horizontal isotherms. Although two-dimensional conductive cooling models show that heat adve ction occurs, the same models suggest the heat advection dissipates rapidly, probably in less than a few million years after the onset of extension (Ketchum, 1996; Carter et al., 2006 ). Thereafter, a st able syn-extensional geothermal gradient is restor ed (Foster and John, 1999). T hus, the complications related to heat advection from initial exhumation of deeper, warmer rocks can be avoided when making detachment fault slip rate estimates by using rock samples with mineral cooling ages at least a few million years younger than the onset of extension (Stockli, 2005). In the case of the ACC, the previously zero aged mica cooling ages appropriate for the slip rate estimates are 2 Ma younger than the onset of extension at ~53 Ma. The potential pitfalls associated with advection of heat in the early st ages of extension are effectively avoided by using these thermochronol ogical data to estimate a slip rate.

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163 Another possible cause for an unstable syn-extensional geothermal gradient, resulting in unrealistic slip rate estimat es, is heating caused by emplacement of voluminous intrusions into the active detachment fault zone (Ehlers, 2005; Stockli, 2005). This is a valid concern and should always be addressed because synkinematic intrusions are common in metamorphic core complexes (e.g., Lister and Bald win, 1993; Foster et al., 2001). Therefore, it is necessary to pr ovide some age constraints on the emplacement of intrusions within the exhumed lower pl ates of metamorphic core complexes when making slip rate estimates using lower plate transect thermochronological data (Stockli, 2005). However, three lines of evidence indi cate the Eocene granitoi d intrusions within the ACC lower plate were emplaced during th e early phases of extension and therefore should not interfere with the slip rate estima tes made here. (1) There is no evidence for radial mica cooling age patterns within or s ounding any of the intr usions (i.e., the mica cooling ages young to the ESE, th e direction of tectonic unroof ing). A radial pattern in cooling ages would be expected if one or mo re of the intrusions were emplaced after the country rock and other intrusions had cooled below mica Ar closure temperatures due to tectonic unroofing (e.g., John and Foster, 1993). (2) The youngest intrusion (as indicated by cross-cutting relationships in the field) porphyritic two-mica granite, yielded mica cooling ages of ~51-48 Ma. Th ese cooling ages indicate that the other intrusions must have been emplaced earlier, probably earlier than ~51-50 Ma. (3) Zircon populations from biotite granodiorite and biotite granite overprinted by greenschist facies mylonitic fabrics within the eastern e xposed mylonite zone yielded 206Pb/238U weighted mean ages 53.0 1.2 and 47.1 0.8 Ma (2 errors), respectively (Foster et al., 2006a). The emplacement of biotite granodiorite is inte rpreted to have occurred at the onset of

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164 extension or just slig htly before. The younger 206Pb/238U weighted mean age for the biotite granite should be considered a mi nimum age of emplacement because several zircons analyzed from this sample gave 206Pb/238U ages of ~53-50 Ma, indicating incorporation of zircon material the age of the granodiorite. Thus, these two intrusions and other intrusions within the ACC lower were most likely emplaced 0-3 Ma after the onset of extension in the ACC. By ~50 Ma the syn-extensional geothermal gradient would have been restabilized. Therefore, real istic slip rates estimates can be made from the lower plate micas that were effectively ze ro aged prior to the onset of extension at ~53 Ma. Slip rate estimates were made using both biotite and muscovite cooling ages separately to compare results from the two different isotopic systems. To make these slip rate estimates, the mica cooling ages and erro rs (as presented in Figure 7-5a, both age and 1 km distance errors) were first imported in to the computer program Isoplot v. 3.09a (see Ludwig,1991) where straight lines were fit to the ther mochronological data using least-squares regressions. In Isoplot, the least-square s regressions (using Isoplot regression mode 1) are made with the original algorithms of York (1969) and the error propagation and correlation between the ther mochronological data points follow that of Titterington and Halliday (1979; see Isoplot us ers manual, p. 21, Ludwig, 2003). Note the biotite cooling age from WG04-033, a sample collected from a aphanitic porphyritic dacite dike, was excluded from the least-s quares regression and s ubsequent slip rate estimate made from the biotite thermochronolog ical data. This dacite dike and several others in the central part of the current study area were emplaced relatively late and their

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165 cooling histories may be significantly differe nt from the country rock and therefore should not be included in the slip rate estimate. Figure 7-6 shows the results from the Isopl ot least-squares regressions made from both the biotite and muscovite thermochronologic al data. For biotite, the regression slope (m) = -1.07 0.34 and for muscovite m = -1.15 0.62 (regression slope errors are 2 ). Slip rate estimates for the detachment were made from these least-squares regressions by taking the inverse of the absolute values of the regression slop es and their errors (e.g., see Foster and John, 1999; Brichau, et al., 2005). The slip rate estimates for the detachment are shown in Figure 6 as well. For biotite, the calculated slip ra te = 0.93 + 0.44 / 0.23 km/Myr (or cm/yr) and for muscovite = 0.87 + 1.09 / 0.31 km/Myr (s lip rate errors are 2 ). The asymmetry of the slip rate errors is derived from adding and then subtracting the regression slope errors from the regression slopes. Despite the larger positive error associated with muscovite slip rate estimate, the closeness the mica derived slip rates indicate the slip rate estimates for the detachment made here are most likely realistic (i.e., the two separate isotopic systems give virtually the same slip rate). The larger positive error of the muscovite slip rate is the result of some scatter in the muscovite cooling ages at ~10-11 km in the slip direction and a relatively large error ( 2.3 Ma) in the mu scovite cooling age from sample DF02-116 at ~19 km, mylonitic two-mica granite sampled fr om upper Clear Creek in the east (Fig. 7-2 5a). Together, the muscovite and biotite slip rate estimates corres pond to an average slip rate = 0.90 + 0.59 / 0.19 km/Myr for the detach ment (a range of 1.5-0.71 km/Myr). It is important to note that these slip rates estim ates are time-averaged, meaning that slip

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166 B A B A Figure 7-6. Slip rate calculations. A) Slip rate calculated from muscovite cooling ages. B) Slip rate calculated from bi otite cooling ages (see text). along the detachment is averaged over the time interval from ~51-39 Ma, the range of cooling ages used in the slip rate estimates. Therefore, these slip rate estimates do not

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167 account for increases or decreases in slip along the detachment during the time period from ~51-39 Ma (Foster a nd John, 1999; Stockli, 2005). Original Detachment Geometries For over two decades now, considerable controversy has been focused on the original geometries of low-angle normal detachme nt faults (i.e., the dip of the fault in the earliest phases of extension) that bound me tamorphic core complexes (see Wernicke, 1992, 1995 or Carney and Janecke, 2005 for a re view). The basis of the a debate is whether these complex-bounding low-angle normal detachment faults originated within the brittle upper crust with shallow ( 30) dips, as many are presently observed in the field, or with steeper ( 45 or even 60) dips; the latter im plies some subsequent tectonic mechanism must alter the original st eep fault geometries. There is a generally consensus, however, among most workers that regardless the dip of fault in the brittle upper crust, normal faults become sub-horizonta l at mid-crustal depths due to interaction with the brittle-ductile transition which acts as a crustal-scale stress guide (e.g., Lister and Davis, 1989). At one end of the debate, Andersonian fault mechanic purists (e.g., Anderson, 1951) argue that normal faults, such as low-angle detachment faults, cannot form originally with shallow dips in the brittle upper crust because the orientation of the maximum principle stress axes (approximately vertical) during extension will not allow such faults to slip. Fault models such as the rolling hinge or domino-style models have been employed to explain the presentday low-angle orienta tion of many complexbounding normal detachment faults (e.g., Spencer, 1984; Wernicke and Axen, 1988; Buck, 1988; Wernicke, 1992; Brady et al., 2000). These two models hold that normal

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168 faults originally form within the brittle upper crust with steep dips and then are rotated to shallower dips later in their evolutions. Th e rolling hinge model calls for the rotation of steeply dipping detachments at their shallo wer crustal reaches by isostatic rebound of the lower plate upon its vertical exhumation. The isostatic rebound is thought to be caused by unroofing of the lower plate itself and in some cases, buoyant forces created from synkinematic intrusions (Wernicke and Axen, 1988; Lister and Baldwi n, 1993). According to the domino-style model, older, steeper nor mal faults become inactive and then rotated to more gentle dips by a younger generation(s) of synthetic normal faults (see Brady, 2000 or Carney and Janecke, 2005 for a review of these models). An overwhelming proportion of current seismic data indicates that active normal faults are slipping with steep dips (>30, Wernicke, 1995) suggesting a steeper dip origin for normal faults in the upper brittle crust is most geol ogically plausible; these seis mic data support the rollinghinge and domino-style models for steep-t hen-shallow dip normal fault evolution. On the other hand, several workers provide strong evidence for a low-angle, or shallow dip ( 30) origin for normal faults in the br ittle upper crust. These workers use thermochronology (see example below in this study), structural reconstructions, and stratigraphic reconstructions to argue for a shallow dip origin of normal detachment faults that bound several major metamorphic core comp lexes in the western United States (e.g., John, 1987; Davis, 1988; Foster et al., 1990; Scott and Li ster, 1992; John and Foster, 1993; Dokka, 1993; Carney and Janecke, 2005). In addition, rela tively recent seismic data document active low-angle normal fau lting in the juvenile metamorphic core complexes of the DEntrecasteaux Islands, Papua New Guinea indicating an original shallow dip origin for low-angl e detachment normal faults is also geologically viable in

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169 nature (Abers, 1991; Abers et al ., 1997; Abers, 2001). The ap parent lack of evidence for shallow dipping active normal faults in many previous seismic studies may simply be because slip on low-angle normal faults is less frequent in most current tectonic environments and as a result less recorded (Abers, 1991; Wernicke, 1995). Therefore at present, it appears that normal detach ment faults bounding many metamorphic core complexes could have originated in the brittle upper crust with steep or with shallow dips as most are now observed in the field. Constraints on the Original Detachment Geometry Apparent mineral cooling ages obtained from rock samples collected across the exhumed lower plates of metamorphic core co mplexes parallel to the direction of tectonic unroofing have also been utilized to estimate the original geometries (i.e., the fault dip) of the bounding detachment normal faults (e.g., Fo ster et al., 1990; John and Foster, 1993; Dokka, 1993; Stockli, 2005). Typically, cooling ages from multiple temperature sensitive mineral systems (e.g., micas, K-feldsp ar, and apatite) are plotted on a single age vs. distance in slip direct ion diagram as the one desc ribed above. Given enough thermochronological data, the pos itions of more than one pal aeoisotherm (e.g., the base of the mica PRZ and apatite partial anneali ng zone, the PAZ) can be identified from inflection points on the diagram (Foster and J ohn, 1999). If the geot hermal gradient at the start of extension is know n or can be estimated then paleodepths to the identified palaeoisotherms can be calculated. The orig inal dip of the detachment normal fault is then simply found by the difference in the pa leodepth between the palaeoisotherms (rise) divided by the distance between th e palaeoisotherms in the direct ion of slip (run). Often, a paleodepth (y-axis) vs. distan ce in slip direction (x-axis) diagram is constructed from this information and the origin al dip of detachment fault is measured directly from the

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170 diagram (e.g., Foster and John, 1999). The orig inal dip of the detachment fault can also be determined by the relations hip: dip (in degrees) = sin-1 (z / h), where z is difference in paleodepth between palaeoisotherms and h is the inferred length of the detachment fault measured along a sloped line connecting the palaeoisotherms on the paleodepth vs. distance in slip direc tion diagram (Dokka, 1993). The above methodology for constraining the original dip of normal detachme nt faults can be used provided that: (1) cooling of the lower plate was due to tectonic exhumation from below a single detachment fault system, (2) any significant la teral variations in the geothermal gradient can be ruled out, (3) two or more well spaced palaeoisotherm can be identified across the exhumed lower plate from the available ther mochronological data, and (4) the geothermal gradient is known or can be calculated for the time the fault dip is to be calculated (Foster and John, 1999). In the case of the ACC, the first two c onditions are met. There is no evidence for more than one generation of normal detach ment faults being responsible for the exhumation and cooling of the lower plate. Seismic imaging of the detachment fault zone just east of the current study area and industry wells reveals a single planar zone mylonite zone dipping to the east at ~20 with no evidence for any major intervening structures (Vejmelek and Smithson, 1995, s ee their Fig. 7-2 2). Also, there is no evidence for a secondary detachment within the lower plate. This is indicated by a relatively smooth change from older to younger cooling ages across the lower plate (besides the change in cooling ages across the base the mica PRZ, Fig. 7-5a). If any intervening structures were present, such as a secondary detachment fault, then an abrupt change in mineral cooling would be seen acr oss the lower plate (e.g., Stockli, 2005). In

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171 addition, the presence of a significant latera l thermal gradient at ~53 Ma is unlikely. Intrusions that were emplaced by ~53 Ma would have ponded below the brittle/ductile transition because this zone acts a barrier to rising magm as (Lister and Baldwin, 1993). The third condition is partially met. The positions of two palaeoisotherms within the ACC lower plate have been defined by mica thermochronological data obtained in this study. These palaeoisotherms are the top and bottom boundaries of the mica PRZ. The temperature of the top of the PRZ fo r biotite and muscovite is ~200C and the temperature of the bottom of the PRZ for biotite and muscovite is ~325C and 375C, respectively. These temperatures were approx imated from the thermal modeling of Lister and Baldwin (1996, see their Figures 9 and 10, p. 97-98). These two palaeoisotherms may be used to estimate the dip of the se gment of the detachment that exhumed the westernmost lower plate. However, another palaeoisotherm(s) from the eastern exposures of the lower plate would help to provide more complete constraints on the detachment geometry. A geothermal gradient at the onset of extension in the ACC (~53 Ma) cannot be estimated from the tilted upper pl ate due of the lack of the necessary upper plate thermochronology. In addition, the A CC upper plate is perv asively shattered by high-angle brittle normal faults making its in appropriate for estimating a paleogeothermal gradient for the ACC at the start of extensi on (e.g., Foster and John, 1999; Stockli, 2005). Despite these apparent short comings, important limits can be placed on the geometry of the detachment fault at ~53 Ma with the thermochronological data obtained in this study. First, the position of anot her paleoisotherm can be approximated in the easternmost exposure of the lowe r plate within the greenschist facies mylonite zone. This paleoisotherm can be used to more completely constrain the original geometry of the

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172 detachment. Because the granitoid mylonites of this mylonite zone formed as a result of ductile deformation at lower to middle greensc hist facies conditions (as indicated by their mineral assemblages and metamorphic textures see descriptions a bove) they could not have been subjected to temperatures great er than ~ 400C (Spear, 1993). In addition, because the micas within these mylonites we re effectively zero aged until ~40-39 Ma they were at least at 350C, and more likel y >375C when extension began at ~53 Ma (Lister and Baldwin, 1996). Therefore, as a conservative approximation, a paleoisotherm representing a temperature of ~400C can be pl aced within the greenschist mylonite zone in the easternmost exposures of the lower pl ate. The ~400C palaeoisotherm was placed at the sample locality of DF02-116a. This sa mple was collected fr om mylonitic two-mica granite at the head of Clear Cr eek structurally high within the greenschist mylonite zone (2573 m). Presumably, the greenschist myl onites at the sample locality of DF02-116a and vicinity were closest to the base of the detachment fa ult plane at 53 Ma. The ~400C paleoisotherm approximated here and those re presenting the top and bottom of the biotite PRZ (at ~200C and ~325C, respectively) are contoured on a simplified geological sketch map of the exhumed ACC lower plate in Figure 7-7. Using the three palaeoisotherms identified across the ACC lower plate and a range of paleogeothermal gradients a set of possible original faults geometries can be calculated for the Anaconda detachment at ~53 Ma. Figur e 7-8 displays a distan ce in slip direction (x-axis) vs. paleodepth (y-axis) diagram where the original geometry of the detachment has been reconstructed using paleogeothermal gradients of 50C/km, 35C/km, 30C/km, and 25C/km. For each paleogeothermal gr adient the paleodepths of samples WG04114, WG04-052, and DF02-120, representing the 200C, 325C, and 400C

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173 paleoisotherms, respectively, were calcul ated simply by dividing the paleoisotherm temperatures by the paleogeothermal gradient s. Lines connecting the palaeoisotherms (and samples locations) for each paleogeotherm al gradient define the shape of the detachment at the onset of extension in the ACC at ~53 Ma. Note the paleoisotherms, as defined by the sample localities of WG 04-114, WG05-052, and DF02-114, must have been positioned structural beneath the detachme nt fault at ~53 Ma (i.e., within the lower plate). Therefore, the actual detachment fa ult plane must have been slightly shallower than as depicted in Figure 7-8. However, be cause erosion is consid ered to be relatively minor across the ACC lower plate the detachment was probably not more than ~1 km above the paleodepths indicated in Figure 7-8. Notably, all four reconstructions reveal a listric shaped (curved) fault geometry for the Anaconda detachment at ~53 Ma (Fig. 7-8). The listric nature of the detachment is defined by a marked break in the fault dip at 5.2 km in the slip direction which corresponds to the ~325C palaeo isotherm, the bottom of the biotite PRZ, and the sample locality of WG04-052. West of (structurally above ) this location, the detachment must have been steep because the minimum difference in temperature between samples WG04-114 and WG04-052 was ~125C and thes e two samples are separated by a distance of only 1.9 km in the slip direction. However, east of (structurally beneath) this location the dip of the detachment must have been much shallower because the difference in temperature between WG04-052 and DF02120 was not more than ~75C and these two samples are separated by a distance of 13.6 km in the sl ip direction (Fig. 7-8). The significant decrease in the apparent dip of the detachment at 5.2 km (the base of the biotite PRZ, WG04-052 sample locality), defining the listric shape of the fault, is

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174 Figure 7-7. Paleoisotherm contour map. Paleoisotherms refer to estimated temper atures of the ACC lower plate directly beneath the detachment at the onset of extension at ~53 Ma. The 200C and 325C paleoisother ms represent the t op and bottom of the biotite partial retention zone for biotite. The 400C isotherm was estimated from the metamorphic grade of the greenschist mylonites in the eastern lower plate (see text).

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175 Figure 7-8. Distance in slip direction vs. paleodep th diagram showing geometries for the detachment at the onset of extension at ~53 Ma. Each reconstruction was made with a different ge othermal gradient to sh ow the range of possible geometries. Regardless of the geot hermal gradient chosen, all the reconstructions indicate the detachment was strongly listric shaped at ~53 Ma with a steep upper crustal portion and a sub-horizontal middle crustal portion (see text). consistent with deeper portion of the fault en tering the brittle-ductile transition (BDT, at ~300-350C) with increasing dept h to the east (Lister and Davi s, 1989). Within the BDT, the yield strength of the brittle upper crust is at a maximum (yield strength increases with depth without substantial pore fl uid pressure present) just above the weaker, warmer, and more ductile lower crust. Therefore, when detachment normal faulting occurs at the BDT, detachments must form with sh allow dips to accommodate the strain

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176 incompatibility between the brittlely deforming upper crus t and plastically deforming lower crust separated by the BDT (Lister and Davis, 1989). The shallowest reconstructed detachment fault dip was obtained with a paleogeothermal gradient of 50C/km (Fig. 7-8) In this fault reconstruction, above the base of the biotite PRZ within the brittle uppe r crust, the reconstruc ted fault dip is 54. Below the base of the PRZ (within the BDT and slightly below), the fault dip is subhorizontal at 7. Note that the 50C/km reconstruction made here is not viable because a geothermal gradient of ~50C/km was in place during upper-amphibolite facies metamorphism (~700-800C) in this part of the lower plate during the late Cretaceous (i.e., within the LISZ, see thermobarometric co nstraints, this study). However, mylonites exhumed along the detachment were deformed under lower to middle greenschist facies conditions (~350-400C maximum temperatures ). Nonetheless, the 50C/km fault reconstruction is presented to show the absolu te minimal dips for the detachment fault in the brittle and ductile parts of the crust. The detachment fault dips obtained from the fault reconstructions using the 35C/km, 30C/km, and 25C/km paleogeothermal gradients are probably more realistic because they are typical of rapidly extending the crust today (e.g., Sags et al ., 200?) and middle Tertiary me tamorphic core complexes in the Colorado River extensional corridor (e.g., Fo ster et al., 1991; Fitz gerald, et al., 1991). These reconstructions show the detachment dipping at 61 and 65 in the upper brittle crust and at 9 and 10 within the BD T and below for the 35C/km and 30C/km paleogeothermal gradients, re spectively (Fig. 7-8). The fa ult geometry reconstruction made with the 25C/km paleogeothermal grad ient give maximum values of dip possible

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177 for the detachment at ~53 Ma. With a pale ogeothermal gradient 25C/km the dip of the detachment is 70 in the brittle upper crust and 12 within the BDT and below. The fault dip reconstructions made here show that the shallower portion of the detachment fault formed in the brittle upper cr ust at ~53 Ma probably with an original dip of ~61-65. A detachment dip in this ra nge for the brittle upp er crust is in good agreement with Andersonian fault mechanic s which states that during extension the maximum principle stress axes ( 1) are oriented at 90 and br ittle normal faults should form with dips >60 (Anderson, 1951). W ithin the BDT and slightly below, however, 1 can become rotated substantially from vertical due the influence of the crustal-scale stress guide created by the high yield strength of the BDT. As a result, detachment faults form with sub-horizontal dips at these depths (Lister and Davis, 1989). Therefore, dips estimates of 9-10 made here for the deeper portion of the detachment fault are certainly reasonable. Magnitude of Offset on the Detachment The magnitude of offset across a given fa ult depends on the fault geometry, slip rate, and duration of slip on th e fault (Scott et al., 1998). Ther efore, the amount of offset (including both vertical and horizontal disp lacement) across a fault can be estimated provided these aspects of the fault can be accu rately estimated. With the constraints on the geometry, slip rate, and duration of slip constraints summarized above the magnitude of offset on the Anaconda detachment can be estimated for the time interval of ~53-40 Ma. In order to estimate the offset ac ross the detachment during this time three assumptions are made: (1) the slip rate on the detachment between ~51-40 Ma can also be applied to slip on the fault between 53-51 Ma, 2) the geometry of the detachment did

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178 not change significantly between ~53-40 Ma, an d 3) the geothermal gradient remained approximately stable from ~53-40 Ma. Figure 7-9 shows a reconstruction diagram for the detachment fault geometry at ~53 Ma with the paleoisotherms represen ting temperatures of 200C, 325C, and 400C assuming a paleogeothermal gradient of 35C/ km. The amount of vertical exhumation and horizontal displacement facilitated by the detachment from ~53-40 Ma can be calculated by moving a point along the fault su rface for 13 Myr at the estimated slip rate and then measuring the displacem ent directly from the diagram. The amount of vertical and horizontal displacement on the detachment has been calculated for the range of possible slip rates on the detach ment between 53-40 Ma (average slip rate = 0.90 + 0.59 / 0.19 km/Myr, 1.49-0.71 km/Myr). If rock X, representing the DF02-120 greenschist facies biotite granite mylonite, is displaced along the detachment at a slip rate of 0.71 km/Myr for 13 Myr, the total amount of ver tical and horizontal displacement = 1.75 km and 8.75 km, respectively. Using the average s lip rate of 0.90 km/Myr the total vertical and horizontal displacement across the detachme nt = 2.0 km and 11.6 km, respectively. Furthermore, if rock X is displaced along the detachment at a slip rate of 1.49 km/Myr then the total vertical and horizontal disp lacement from ~53-40 Ma = 7.3 km and 16.1 km, respectively (Fig. 7-9). These calculations indicate the possible range of vertical and horizontal displacement on the detachment from ~53-40 Ma is 1.75-7.3 km and 8.75-16.1 km, respectively. The range of possible displa cement along the detachment during this time interval is large because of the large errors on the slip rate estimate. However, because the greenschist facies mylonites cooled rapidly ( 125C/Ma) they had probably just

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179 passed upwards into the steeper portion of the detachment by ~40 Ma. This would require a minimum of 2.4 km and 13.8 km of vertical and horizontal displacement, respectively, on the deta chment from ~53-40 Ma. Note the constraints on the magnitude of offset on the Anaconda detachment made here are only for vertical and horizontal di splacement of the now exhumed greenschist mylonites during the time interval of ~5340 Ma. During this time, the greenschist mylonites were uplifted to crustal depths above the brittle-ductil e transition. These constraints do not take into account the con tinued vertical and horiz ontal displacement of the mylonites into the brittl e upper-crust and towards the surface, presumably from ~4030 Ma (see constraints on the dur ation of extension above). Another way to constrain the amount of horizontal displacement on the detachment, which does take into account extension from ~53-30 Ma is to reconstruc t the upper plate to the lower plate using a structural pinning point. Figure 7-10 displays a geologic map of the ACC showing the Anaconda detachment, which separates the lowe r plate (to the west) from the upper plate (to the east). Two intrusions are highlighted here: 1) the SLS granodiorite exposed in the ACC lower plate labeled A and 2) a detached granodior ite block in the upper plate labeled B. These two intrusions share ve ry similar mineralogies and textures. In addition, a sample from the SLS granodior ite (WG04-114) and the detached upper plate granodiorite (DF02-114) gave very similar biotite 40Ar/39Ar cooling ages (74.0 0.9 and 76.3 1.1 Ma, respectively, Table 6-2). Furthermore, the major and trace element compositions of these two samples suggest th e two intrusions are related by fractional crystallization and/or assim ilation (AFC, Appendix D). Th ese similarities suggest the two intrusions were originally part of the same intrusive body prior to the onset of

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180 extension in the ACC at ~53 Ma. The fact that the detached u pper plate granodiorite block lies to the ESE of the SLS granodior ite (the approximate extension direction, Chapter 3) supports its origin as part of the SLS granodiorit e as well. Reconstructing the SLS granodiorite and detached upper plate granodiorite indicates ~25-28 km of horizontal displacement was facilitated on the Anaconda detachment from ~53-30 Ma (Fig. 7-10). Note if the-the Anaconda det achment slipped at the average slip rate calculated above (0.9 km/Myr) from ~53-30 Ma then the total disp lacement would be ~21 km. This estimate of horizontal displa cement is similar to the one made using the two granodiorite structural pinning points. The amount of vertical displacement (total exhumation) facilitated by the Anaconda detachment may be better constrained using th e thermobarometric data obtained from the LISZ in this study. Because the LISZ is Late Cretaceous in age (~75-74 Ma, see above and Chapter 4) and it lies st ructurally beneath the Eocene greenschist facies mylonite zone in the eastern ACC (Chapter 3) the pr essure-temperature estimate from the LISZ can be used to constrain the maximum amount of exhumation. The pressure-temperature estimate for sample ME-231, meta-Greyson pelitic paragneiss from the LISZ, gave upper-amphibolite facies pressures and temp eratures of ~3.2-5.3 kbar and 750-825C (Chapter 5). These pressures correspond to crustal depths of ~10-16 km, assuming an increase in pressure of 0.33/km depth. Therefore, a maximum of 16 km of vertical displacement (total exhumation) was facilita ted by the Anaconda detachment after ~53 Ma, assuming no extension occurred between the late Cretaceous (~75-74 Ma) and middle Eocene time. The maximum total ver tical displacement estimate made here (based on thermobarometry) is consistent wi th the detachment reconstruction shown in

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181 figure 7-9. Based on this detachment rec onstruction (using a geot hermal gradient = 35C/km) the greenschist mylonites now e xposed in the eastern ACC were exhumed from a crustal depth of ~13 km, n ear the brittle-ductile transition. Figure 7-9. Distance in slip direction vs. paleodepth di agram showing magnitude of offset on the detachment with variable slip rate from ~53-40 Ma with a geothermal gradient of 35C/km. Ro ck X represents the greenschist mylonites now exposed at the surface in the eastern ACC lower plate (see text).

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182 Figure 7-10. Geologic map of the ACC show ing structural pinning points used to cons train the maximum amount of displacement facilitated by the Anaconda detachment. The two pinning poin ts are (A) the Storm Lake St ock granodiorite in the lower plate and (B) a detached granodiorite bloc k in the upper plate. Reconstructing the detached granodiorite in the upper plate with the SLS in the lower plate indicates ~25-28 km of displacement on the detachment (Modified from Lewis, 1998).

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183 Chapter 8 CONCLUSIONS Age and Pressure-temperature Constraints on High Grade Metamorphism in the ACC lower plate: Origin of the Lake of the Isle Shear Zone Thermobarometric data obtained from the attenuated Lower Belt Greyson migmatitic paragneiss, within the Lake of th e Isle shear zone (LISZ) in the ACC lower plate, show that anatexis occurred at uppe rmost-amphibolite facies pressure-temperatures conditions at ~3.2-5.3 kbar and 750-825C (Cha pter 5). Granitic leucosome found within pressure shadows between boudins shows that ductile deformation accompanied this metamorphism (Chapter 3). U-Pb zircon crys tallization ages from the deformed quartz diorite sill in the LISZ and the Storm Lake Stock (SLS) granodiorite which cross-cuts the LISZ indicate that uppermost-amphibolite f acies metamorphism and ductile deformation in the LISZ occurred at ~75-74 Ma (C hapter 4). The meta-Greyson migmatitic paragneiss, located directly beneath the greensc hist mylonite zone a nd brittle detachment in the ACC, are not the result of decomp ressional anatexis duri ng exhumation of the ACC lower plate in the Eocene (as in the BCC to the west, e.g., House et al., 1997; Foster et al., 2001). Rather, the me ta-Greyson migmatites were fo rmed during anatexis related to the emplacement of the quartz diorite sill in the Late Cretaceous in the LISZ, most likely in a compressional tectonic setting. Exhumation and Cooling History of the Middle Eocene Anaconda Metamorphic Core Complex: The thermochronological data set obtained from the Anaconda metamorphic core complex (ACC) lower plate transect in this study provides severa l constraints on the

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184 exhumation and cooling history of the ACC dur ing the Eocene: (1) The age of the onset of extension in the ACC is indicated to be ~53 Ma by the marked break in the slope of the cooling age curve on the age vs. distan ce diagram constructed from the lower plate mica cooling ages (Fig. 7-5). This thermo chronology-based age constraint is in good agreement with the one U-Pb zircon crystall ization age of 53 0.6 Ma from mylonitic granodiorite within the greenschist facies my lonitic shear zone (Foster et al., 2006a). Thus, the onset of extension in the ACC is now well constrained and confirmed to be at ~53 Ma, coincident with the onset of extens ion in the BCC to the west (Foster et al., 2001, 2006a). Furthermore, (2) the cooling ag es from micas in the greenschist facies mylonites in the eastern study area show that extension in the ACC continued until at least ~40-39 Ma. These cooling ages record the time at which ductile deformation ceased in the exposed part of the greenschist mylonitic shear zone and do not take into account their continued uplift into the brittle upper crust. However, apatite fission track ages from the southern Anaconda-Pintlar Range (Foster and Raza, 2002), suggest brittle extension in the ACC continued to ~30 Ma. Thus, extension in the ACC likely spanned from ~53-30 Ma, the synchronous with extensi on in the BCC to the west (Foster et al., 2001; Foster and Raza, 2002). (3) The late ral cooling age gradient, defined by mica cooling ages that progressively young to the east across the ACC lower plate, confirms that tectonic unroofing of the lower plate occurred by top-to-the-east-southeast directed detachment of the upper plate from ~53-39 Ma ; this is consistent with the extension direction suggested by the mineral stretching lineations in the greenschist mylonites (102108, Kalakay et al., 2003). (4) Thermochronol ogical data from the lower plate show the original Anaconda detachment (at ~53 Ma ) was an east-dipping listric-shaped normal

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185 fault with a steeply dipping (54-70) portion in the brittle upper crust and a low-angle (~7-12), sub-horizontal portion in the middle crust near th e brittle-ductile transition (Chapter 7). The ACC lower plate was e xhumed along a single east-dipping mylonitic shear zone and brittle detachment system dur ing the Eocene synthetic to the east-dipping detachment in the BCC to the west. The origin al geometry of the Bitterroot detachment cannot be constrained to such detail due to deeper erosion of the detachment zone (D. Foster, per comm.). (5) Thermochronological data obtained in this study were also used to constrain the slip rate on the ACC detachment. These da ta show that between ~53-40 Ma, the timeaveraged slip rate on the detachment was ~0.9 km/Myr. Currently, slip rates have not been constrained for detachments in any other metamorphic core complex north of the Snake River Plain (SRP). Thus, no comparison can be made with these core complexes, including the BCC at this time. However, the slip rate for the ACC detachment is significantly low compared to detachment time -averaged slip rates of ~3.0-7.0 km/Myr from middle Tertiary metamorphic core comp lexes in the Colorado extensional corridor of the southwestern United States (e.g., Fost er and John, 1999). Therefore, the rates of extension in middle Eocene metamorphic core complexes north of the (SRP) may have been much slower, at least initially, compared to later extension in the middle Tertiary metamorphic core complexes to the south. To fully compare the rates of extension in these separate regions slip rates are needed for later brittle extens ion in the ACC (i.e., during the ~40-30 Ma time interval) and slip rates are needed for the other core complexes north of the SRP. Low temperat ure thermochronological data (e.g., apatite

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186 fission-track and (U-Th)/He apatite cooling ages) can be used to constrain this later lower temperature brittle extens ion (e.g., Stockli, 2005). (6) Constraints have been placed on the amount of offset on the ACC detachment using the thermochronological da ta set obtained in this study. These data show that a minimum of 13.8 km and 2.4 km of horizontal displacement and vertical exhumation, respectively, occurred along the detachment between ~53-40 Ma. However, this amount of offset on would have exhumed the greenschi st mylonites to crustal depths only just above the brittle-ductile transition which was probably at a depth of ~10-12 km (assuming a geothermal gradient of 35C/km, Fi g. 7-9). These estimates of offset do not account for uplift of the mylonite s in the brittle u pper crust, presumably with continued brittle extension and exhumation of the green schist mylonites thr ough the shallower crust from ~40-30 Ma. Major and trace element geochemistry from similar Late Cretaceous granodiorite intrusions from the lower plate a nd the detached upper pl ate suggest they are related by fractional crystallization (with mi nor assimilation, AFC, Appendix D) and originally part of the same pl uton (the SLS). Reconstructing these granodiorite intrusions shows the ACC detachment accommodated ~25-28 km of displacement after the onset of extension at ~53 Ma. Interestingly, if a s lip rate of ~0.9 km/M.y. is assumed for the Anaconda detachment from ~53-30 Ma then ~22 km offset was accommodated along the detachment during this time interval. This is similar to the displacement estimate made by reconstructing the granodi orites from the lower and upper plates indicating a displacement estimate on the order of ~22-28 km is realistic. (7) Thermobarometric data from the LI SZ show that upper amphibolite facies metamorphism in the Late Cretaceous occurr ed at ~3.2-5.3 kbar, which corresponds to

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187 depths of ~10-16 km depths (0.33 kbar/km) fo r the shear zone at ~75-74 Ma (see above). Therefore, a minimum of ~10 km of vertical exhumation was facilitated along the ACC detachment during the Eocene. This estimat e is consistent with the exposed Eocene mylonites being of lower to middle greenschist facies and located ne ar the brittle-ductile transition (at a depth of ~10-12 km, assuming a 35C/km geothermal gradient) at the onset of extension. The conclusions summarized above show several aspects of Eocene extension in the ACC are remarkably similar to extension in the BCC located to the west. However, the data obtained in this study show important differences between these core complexes which include: (1) Extension in the ACC was accompanied by greenschist facies metamorphism and not by upper-amphibolite facies metamorphism and decompression anatexis as in the BCC. (2) Therefore, total (vertical) exhum ation in the ACC ( 10-16 km) is much less than in the BCC, whic h is estimated to be ~20-25 based on thermobarometric data from early to middle Eocene migmatites in the eastern lower plate. In addition, (3) Di splacement in the ACC (~22-28 km) is approximately half the displacement in the BCC, which is ~40-50 km based on based on reconstructions of Cretaceous dioritic plutons in the detached uppe r plate with similar aged dioritic plutons in western BCC lower plat e (Foster et al., 2006a). The similarities (and differences) in th e timing and kinematics of extension between the ACC and BCC permit the proposed tectonic model for Eocene extension south of the Lewis and Clark line (e.g., O Neill et al., 2004). These two metamorphic core complex therefore represent one continuous and integrated extensional system which accommodated large-scale crustal extension (~ 60-75 km of eastward displacement) south

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188 of the Lewis and Clark line from ~53-30 Ma. Both of the metamorphic core complexes were exhumed along single separate, but s ynthetic, east-dipping detachments. In addition, the thermochronological from this study show the Anaconda detachment was sub-horizontal near the brittleductile transition and therefor e does not likely merge with the Bitterroot detachment at depth. Ther efore, a nested position for the Anaconda detachment with respect to the Bitterroot detachment is geologically plausible. Regional Tectonic Context The results from this study show that synchronous exhumation of the BCC and ACC accommodated large-scale extension sout h of the LCL beginning in the Eocene at ~53 Ma. The onset of extension in the BCC and ACC was coincident with the onset of extension in metamorphic core complexes nor th of the LCL in northeastern Idaho and eastern Washington (the Priest River, Kettle, and Okanagan), southern British Columbia (the Shuswap) and within the LCL (the Clear water). Extension in all these metamorphic core complexes began during the time interv al of ~54-52 Ma (Fos ter et al., 2006a, and references cited therein). These metamorphi c core complexes formed in the previously thickest parts of the Sevier hinterland immediately following the end crustal shortening in the foreland fold-and-thrust belt directly to the east. The early extension in these metamorphic core complexes was accompanied by voluminous back-arc KamloopsColville-Challis-Absaroka (KCCA) magmatism; this magmatism was concentrated in regions where the Eocene extension was most intense (Armstrong and Ward, 1992; Morris et al., 2000; Breitsprecher et al., 2003; Foster et al., 2001). Exhumation of metamorphic core complexes in southern British Columbia and the northwestern United States was linked region al dextral transtension along major dextral strike-slip fault zones during th e Eocene. In southern Britis h Columbia, extension in the

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189 Shuswap metamorphic core complex was linke d to dextral strike-slip motion along the approximately margin-parallel northern Ro cky Mountain trench and Yalakom-Ross Lake fault zones (Price and Carmichael, 1986; Fost er et al., 2006a). In the northwestern United States, regional dextral transtension along the northwest-wes t trending Lewis and Clark line drove extension and exhumation of metamorphic core complexes during the Eocene (Foster et al., 2006a). Eocene extension in metamorphic core complexes to the north of the LCL was directed approximat ely E-W (e.g., Shuswap metamorphic core complex) and ENE-WSW (e.g., Priest River metamorphic core complex). South of the LCL, Eocene extension was directed to th e ESE in the BCC and ACC, approximately parallel to the strike of the LCL. Therefor e, the LCL apparently served as a regional accommodation zone between E-W and ENE-WSW di rected extension to the north of the line and ESE directed extension to the sout h within the BCC and ACC (Foster et al., 2006a). Regional dextral transtension, exhumation of metamorphic core complexes, and large-scale crustal extension along major stri ke-slip faults zones in southern British Columbia and the northwestern United States during the Eocene was most likely driven by changes in plate boundary forces along th e western margin of the North American Plate. Plate reconstructions show the angl e of convergence between the North American and Pacific Ocean basin plates became incr easingly oblique beginning in the early to middle Eocene, coincident onset of larg e-scale extension and exhumation of the metamorphism core complexes southern Bri tish Columbia and the northwestern United States (Haeussler et al ., 2003). An increase in the obli quely of convergence (increasing right-lateral convergence with time) woul d have created traction along the western

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190 margin of the North American Plate driving dextral transten sion and extension inboard of the plate boundary during the Eocene (Breitsprecher et al., 2003). Plate reconstructions also indicate the presence of a slab window (a gap between subducting plates) beneath the North American Cordillera of southern British Columbia and the northwestern United States during the early to middle Eocene (Thorkelson and Taylor and 1989; Breitsprecher et al., 2003; Haeussler et al., 2003). The approximate orthogonal subduction of the Kula-Farallon ri dge (or the Resurrec tion-Farallon ridge, e.g., Haussler et al., 2003) during the Eocene would have created the slab window. The presence of a slab window beneath the North American Cordillera of southern British Columbia and the northwestern United States at this time is consistent with voluminous Eocene KCCA magmatism in the back arc regi on (Breitsprecher et al., 2003). The backarc KCCA magmatism may have been the result asthenosphere upwelling through the slab window and subsequent partial melting of the lithosphere (Foster et al., 2006a). Heating and partial melting of the lithosphere above a slab window may have caused thermal weakening of the lithosphere which may have also helped to facilitate large-scale extension and exhumation of me tamorphic core complexes in southern British Columbia and the northwestern United States be ginning in the early to middle Eocene (Breitsprecher et al., 2003).

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191 APPENDIX A METHODOLOGIES Field Mapping and Sampling Methods Prior to this study a detailed, accurate ma p of the geology from Storm Lake to the Mill Creek area of the Anaconda-Pintlar Range was not available. Thus, production of such a map was necessary to work out the co mplex structures exposed in the ACC lower plate. Mapping these structures in de tail was essential in understanding the deformational history of the lower plate rocks. Furthermore, this mapping was needed to determine spatial relationships between rock samples collected for the analytical work of this study (i.e., 40Ar/39Ar thermochronology, U-Pb geochronology, thermobarometry, and major and trace element geochemistry and modeling). The geology of the ACC lower plate e xposed east of Storm Lake was mapped during multiple transects made by foot over a five week a nd one week period during the summers of 2004 and 2005, respectively. During the transects, unit contacts were hand drawn on several 1:15000 scale t opographic base maps created from the Storm Lake, Mt. Evans, and Mt. Haggin 7.5 minute quadrangles. In addition, one hundr ed structural data points were collected along the mapping tr ansects. These include strike-and-dip measurements on metamorphic foliations, unit contacts, fault surfaces, joints, and cleavage. Lineation measurements made included mesoscopic fold hinges, slickenlines on fault surfaces, and limited stretching or mi neral lineations on foliation surfaces. The hand drawn geologic maps were then digita lly drafted into one continuous 1:24000 scale geologic map using Adobe Illustrator vers ion 10. The mapping of Heise (1983) was

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192 added to the easternmost map area in the Mount. Haggin area. In addition, some mapping of Emmons and Calkins (1914) and Kalakay et al. (unpublished mapping) were added to fill gaps in the final map. The rock samples analyzed in this study we re collected during tr ansects made while mapping in the summers of 2004 and 2005 with the exception of samples ME-6, ME-231, SL-38, Ug-1, DF02-113, DF02-114, DF02-116a, DF02-117a, DF02-118a and b, and DF02-120. Samples ME-6, ME-231, SL-38 and Ug-1 were collected by Tom Kalakay during the summer of 2002. DF02 samples were collected by David Foster in summer 2002. 40Ar/39Ar Thermochronology Methods Sample Preparation and Irradiation In preparation for 40Ar/39Ar thermochronology, thin sec tions were made from select rock samples collected during the summer of 2004. Billets were cut from these samples and sent to Petrographic International for the th in section preparation. The thin sections were necessary to inspect the qua lity of mineral phases in the rock samples to be used in 40Ar/39Ar analyses. Only rock samples which exhibited fresh and unaltered micas, kfeldspars and hornblendes were used for the analyses 40Ar/39Ar analyses. The selected samples were then crushed and milled into a medium to fine-grained size using a Sturtevant rock Jaw Crusher and Bico Pulverizer type UA disk mill, respectively. Each disk milled samples was th en dry sieved in a stack of (from top down) 30, 40, 50, 60, 80 and 100 mesh sieves. Th e 40 and 50 mesh fractions were kept for further separation of biotite, muscovite and k-feldspar. However, in the case of hornblende, 60 and 80 mesh fracti ons were kept for further separation. This was done to avoid inclusion-rich hornblendes in some rock samples as observed in thin section.

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193 The sieved rock samples were then pro cessed with Tetrabromoethane (TBE) and Methylene Iodine (MI) heavy liquids (densities of 2.96 and 3.33 g/cm3, respectively) to separate the micas, k-feldspars and hornblende s based on densities. These separates were rinsed with ethanol (for TBE) and acetone (for MI) 2-3 times following heavy liquid separation. A Frantz magnetic separator Model L-1 was then used to separate the biotite and hornblende from non-magnetic phases. Only the most magnetic biotite and hornblende Frantz separates we re kept to avoid inclusionrich minerals. All mineral separates were then hand picked under a binoc ular microscope for better refinement using standard picking tools (i .e., nylon brushes, Pyrex glass dishes and wax weighing paper). Some hand-picked biotite se parates were given an ultrasonic bath in dionized H2O for approximately fifteen minutes to rem ove any altered materials. The cleansed, hand-picked mineral se parates along with GA1550 biotite flux monitors (98.79 0.5 Ma, see Reene et al., 19 98) were then individually packaged in aluminum foil (~ 10 mg for mineral separate 1 mg for flux monitors) and sealed in a quartz glass tube by melting the ends of th e tube. The mineral separates and flux monitors were irradiated for 2 hrs in a 1.1 MW TRIGA MARK II research nuclear reactor at the Oregon State University Radiat ion Center. For a more detailed description of these facilities and irradiation methods see http://ne.oregonstate.edu/f acilities/radiation_center/ Note micas, k-feldspar and hornblende from the DF02, Ug-1 and SL-38 samples underwent the same sample preparation and irradiation as described above. However, these samples were irradiated in an earlier irradiation batch for ten hours rather than two and no thin sections were made available.

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194 40Ar/39Ar Analytical Instrume ntation and Procedures The 40Ar/39Ar analyses were carried out in the noble gas laboratory at the Department of Geological Sciences, University of Florida. A combination of both laser ablation and furnace step-heating techniques were utilized to extract Ar gas from mineral separates. Micas were step-heated by both laser ablation and furnace, where k-feldspar and hornblende separates underwent furnace st ep-heating only. Laser ablation stepheating of mica separates was facilitated by a water-cooled New Wave Research model MIR10 30W CO2 laser. During laser step-heating the New Wave laser was manually controlled using LAS (laser ab lation software) version 1.3. 0.1 by New Wave Research. Mica separates were ablated a total of 5-15 step s under a 1750 m continuous wavelength focused laser beam at 2-5.5% power. A fina l laser fusing-step was employed at 10-12% power. The step-heating schedule used fo r laser ablation varied and was adjusted accordingly to maximize Ar gas output for the mineral separates. A water-cooled doublevacuum resistively heated furnace was used to step-heat k-feldspar and hornblende separates. The furnace step-heating analys es were controlled manually using LabSpec version 3.1 by LabView. Ar gas extracted from laser and furnace step-heating was transferred by vacuumed lines to a getters trap for 5-10 min to re move volatiles. The cleansed Ar gas was subsequently analyzed in a gas-sourced Mass Analyzer Products Model 215-50 mass spectrometer equipped a filament for ga s ionization and a magnetic sector mass discriminator followed by a single faraday colle ctor to measure the isotopic abundances of 36Ar, 37Ar, 38Ar*, 39Ar, and 40Ar. Cold laser blanks were typically analyzed at the beginning of each analytical session and every five steps after. The laser blanks were

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195 made by closing the cleansed laser chamber fo r two min and then passing the blank gas to the mass spectrometer just as a sample. Heated furnace blanks were made by closing and heating the empty furnace to 750C, 1000C a nd 1100C prior to sample step-heating. Data files produced from the argon gas and blan k analyses were then imported into the program ArArCALC version 2.2 by Koppe rs (2002) for data regression and 40Ar/39Ar cooling age calculations. ArArCALC uses Excel by Microsoft to plot argon data tables, age plateaus and isochrons. ArArCALC was al so used to calculate J-values from total fusion analyses of the GA1550 biotite flux monito rs. These J-values were then applied to 40Ar/39Ar cooling age calculations. Thermobarometry In preparation for thermobarometry, two po lished thin sections were prepared from ME-231, migmatitic pelitic cord ierite gneiss sampled from the central LISZ. These thin sections were then carbon coated using a PAC-1 PELCO carbon coater Model 9500. The elemental oxide weight percent data need ed for the ME-231 thermobarometry were acquired using by electron microprobe analyses at the Department of Earth Sciences, Florida International University (FIU). The polished and coated thin sections were then analyzed at FIU using a JOEL Model 8900R Superprobe electron microprobe. The JOEL Superprobe uses five two-crystal wa velength-dispersion spectrometry (WDS) spectrometers and one electron-dispersion sp ectrometry (EDS) detector to measure mineral compositions and for back-scatter imaging, respectively. Before these analyses, the JOEL Superprobe was calib rated using in-house minera l standards to measure Na, Mg, Al, Si, K, Ti, Ca, Cr, Mn and Fe compositions in ME-231 mineral phase. For a list of the FIU in house mineral standards see http://www.fiu.edu/~emlab/Inst_EPMA_standards.html Note these analyses were

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196 carried out with the JOEL Superprobe with the beam current set to 20 nA and an accelerating voltage of 15 kV. The compositions of garnet, biotite, plagio clase and cordierite were measured in order to calculate phase equi libria and average PT estimates for the ME-231 mineral assemblage. Care was taken only to analy ze the most euhedral and clearest mineral grains. Back-scatter imaging allowed fractures and suspicious areas to be avoided. Rimcore-rim spot transects were made across larger intact positions of garnet and cordierite porphyroblasts. In addition, several fine-gra ined biotite and plag ioclase grains were analyzed from the ME-231 matrix. Inclusi ons were rare in garnet and cordierite porphyroblasts and only a few small biot ite inclusions were measured. The selected elemental oxide weight percents from the electron microprobe analyses were subsequently imported into the program AX by Ho lland and Powell (2000) to calculation of cation unit formulas (with ferric iron estimations) and mineral endmember activities. Mineral end member activities were then imported into the program THERMOCALC version 3.2 by Powell et al. (1998) for average PT calculations for the mineral assemblage of ME-231. Pressure-tempe rature calculations were also made using the program Geothermobarometry (GTB) v 2. 1 by Spear and Kohn (1999). The details of thermodynamic calculations made using th e programs AX, THERMOCALC, and GTB are described in the thermobarometry results section. U-Pb Geochronology Sample Preparation In preparation for zircon U-Pb analyses several kilograms of samples WG04-114, Ug-1 and WG05-02 were first crushed to gravel -sized particles using a Sturtevant rock Jaw Crusher. This rock material was then passed through a Bico Pulverizer type UA disk

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197 mill several times until all material fell through a dry 425 m (40 mesh size) sieve. The < 40 mesh rock fractions were then passed ove r an inclined water table twice for density separation. The heavy water table fractions underwent further density separation using Tetrabromoethane (TBE) and Methylene Iodine (MI) heavy liquids (densities of 2.96 and 3.33 g/cm3, respectively). The MI sink fractions were kept and run through a Frantz isodynamic magnetic separator model L-1 to remove the remaining unwanted magnetic minerals. Each fraction was passed through the Frantz separator several times at a setting of 1.5 volts. For each run, the Frantz trough tilt was decreased until the least magnetic fraction of zircons was obtained (at 0-1). Individual zircons were subsequently handpicked from the non-magnetic Frantz fractions using a binocular microscope and standard picking tools (i.e., nylon brushes, Pyrex glass dishes and wax weighing paper). Care was taken to pick the clearest, most euhedral and inclusion-free zi rcons. In addition, smaller zircons (< 300 m) were targeted ove r larger ones to avoid possible inherited cores that if dated would not represent crystallization ages of interest. In preparation for laser ablation, the indi vidual hand-picked zi rcon fractions were mounted on separate pieces of adhesive Buehler Meta-Kleer mo ld base along with several large fragment of FC-1 zircons, the ~1100 Ma extern al zircon standard chosen for the U-Pb analyses (in house 207Pb / 206Pb age = 1086.9 5.3, 207Pb / 235U age = 1091.5 13.4, and 206Pb / 238U age = 1096.7 21.7 Ma). Each unknown and FC-1 pair was then surrounded by a 2.5 cm diameter plastic ring mold and a 5-to-1 mixture of Buehler Epoxicure resin and hardener was poured ove r the zircons. After hardening, the epoxy zircon mounts were ground down to an even flat surface and wet-polished using a Buehler Ecomet 6 variable speed grinder and polisher with 600, 800 and then 1200 grit

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198 disks. Each zircon mount was ground and polis hed to expose an equatorial section of the zircons for laser ablation. The polished zi rcon mounts were cleaned using sterile Chemwipes and 2-proponol just prior to load ing into the laser ablation chamber. LA-MC-ICP-MS Instrumentation Zircon U-Pb isotopic data were acquired in this study using laser ablation multicollector inductively coupled plasma mass spectrometry (LA-MC-ICP-MS). The mass spectrometer employed is a Nu Plasma multi-collector plasma source mass spectrometer produced by Nu-Instruments installed on site at the Department of Geological Sciences, University of Florida. The Nu Plasma ma ss spectrometer is confi gured in a standard Nier-Johnson geometry consisting of a plasma torch with focal lens (cones) followed by a 35 cm radius electrostatic analyzer (ESA), a 25 cm radius laminated electromagnet and variable dispersion zoom lenses. These devices are followed by a detector array comprised of 12 sapphire/ceramic Faraday collectors and 3 ion counters mounted on a single fixed mop plate with electron suppressors. The i on counters are comprised of three discrete dynode electron multipliers and a retardation filter to measure isotopic abundances with higher sensitivity than Fa raday collectors (for more details see http://www.nu-ins.com/npdetail.html ). The Nu Plasma mass spectrometer is configured to measure 238U and 235U abundances on Faraday collector s EX-H and H6, respectively and 207Pb, 206Pb and 204Pb abundances on ion counters IC0, IC1, and IC2, respectively. The Nu Plasma mass spectrometer is coupled to a Nu-Instruments DSN 100 desolvation nebuliser. In this study, the DSN 100 was us ed to introduce a U-Pb calibration solution to the mass spectrometer prior to the analyses (see below). A more detailed description of this Nu Instrument equipment is provided by Simonetti et al. (2005). The Nu Plasma

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199 mass spectrometer has also been coupled w ith a New Wave 213 nm ultraviolet laser manufactured by Merchantek Products for abla ting the solid geologic samples in this study. LA-MC-ICP-MS Analytical Protocol All laser ablations in this study were carried out in the presence of a helium carrier gas. The helium carrier gas was introduced directly into the ablation chamber using external Tygon silicon tubing where it mixes with the ablated sample and aids in its transportation to the plasma torch. Use of helium as a carrier gas during single-spot laser ablation analyses has been shown to reduce within-run Pb-U elemental fractionation and increase instrument sensitivity, both desired results (Simonetti et al ., 2005). In addition, argon gas was subsequently mixed with the heli um and ablated sample mixture just prior to the plasma torch using Tygon tubing and a plastic Y-connect or. The argon gas aids in the sample transportation and reacts with a radio frequency coil in the plasma torch box to produce the high temperature plasma needed to ionize the ablated sample (Thomas, 2001). Prior to each analytical session the Nu Plasma ion counters were calibrated to ensure accurate isotope abundance measurements and proper sensitivity. To achieve this calibration a solution of SRM 3164 U (25 ppb) and SRM 981 Pb (1 ppb) was aspirated into the Nu Plasma mass spectrometer using the DSN 100. The abundances of 204Pb, 206Pb and 207Pb were then measured on the ion counters for a minimum of 1 hour to ensure proper calibration and to warm up the Nu Plasma mass spectrometer. After the ion counter calibration the DSN 100 was disconnected from the Nu Plasma mass spectrometer.

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200 Prior to data acquisition the helium and argon gas flows, laser settings, and NuPlasma instrument parameters were carefully tuned to maximize the signal intensities. A standard zircon and an unknown zircon were ablated separately for ~ 60 s while these parameters were adjusted to obtain the best se ttings. In every analyt ical session, the best results were obtained with the helium and ar gon gas flow rates set to 0.6-0.7 L/min and 1.05 L/min, respectively. In addition, best results were obt ained with a laser setting of 50-60% power, a 2-4 Hz pulse frequency and a laser spot size of 30-60 m. Note that all the standard and unknown zircons were ablate d under the same gas and laser settings throughout each analytical se ssion as determined at the beginning of the session. Isotopic data were acquired during the an alyses using Time Resolved Analysis (TRA) software provided by Nu-Instruments. Before the ablation of each zircon a 30 s on-peak zero was determined on the blank He and Ar gases with clos ed laser shutter. This zero is used for on-line correction for isobaric interferences, particularly from 204Hg which is largely derived from the argon gas. Following blank acquisitions individual zircons underwent spot analyses for ~30-60 seconds. The analyses of unknown zircons were bracketed by analyzing an FC-1 sta ndard zircon for every 2-6 unknown zircons. These standard zircon analyses were later used to correct raw unknown zircon isotopic analyses (see below). An important feature of the TRA is that it allows the analyst to choose the desired portion (or time) from each analysis used in the subsequent isotopic abundance and ratio calculations. This feature a llows the omission of 1) slight spikes in 204Pb over the first few seconds during some of the analyses and 2) Pb-U elemental fractional near the end of each ablation rela ted to deepening of the ablation pit (see Kosler and Sylvester, 2003).

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201 U-Pb Data Reduction and Common Lead Correction Raw isotopic data obtained from the LA -MC-ICP-MS analyses and TRA software were imported into a Microsoft Excel spreadsheet where they were corrected for instrumental drift (change in instrument sensit ivity with time) and isotopic fractionation. The raw isotopic data were corrected for instrumental drift using the formula: R(sam)meas x (1-Corfac x (P/N)) (A-1) where R(sam)meas is the raw isotopic ratio for the sample and Corfac is a correction factor derived from taking the difference between the measured isotopic ratios of the two standards which bracket the measured is otopic ratios of the unknowns (standard 2 minus standard 1) and then dividing this valu e by the measured isotopic ratio of the first standard (standard 1). P is the number of the bracketed unknowns and N is the number of intervals between the unknowns and bracketi ng standards (e.g., five unknowns bracketed between two standards would be numbered 1, 2, 3, 4, and 5 and have 6 intervals). This equation was applied to each set of unknown bracketed between pairs of standards analyses for a given analytical session. The drift corrected unknown zircon isotopic ratios were then corrected for isotopic fractionation using the following equation: R(sam)true = R(sam)meas x [R(std)true / R(std)meas] (A-2) where R (sam)true is the drift and fractionation corrected unknown ratio and R(sam)meas and R(std)meas are the raw isotopic ratios of the unknown and standard zircons, respectively. R(std)meas is the raw isotopic ratio of the first standard measured during the analytical session. R(std)true is the true isotopic ratio of the external standard zircon determined by a separate isotopic analyses; the R(std)true values for the standard FC-1 zircons have been determined by ID-TIMS analyses (207Pb / 206Pb = 0.0762, 207Pb / 235U =

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202 1.9428 and 206Pb / 238U = 0.1850, Paces and Miller, 1993). This standard-samplestandard correction for drift and isotopic fractionation is different from the Tl-doping correction employed in many previous LA-MCICP-MS studies. Recent work has shown this method is equally effective given the exte rnal standard is well characterized, as is FC-1 zircon standard (e.g., Paces and Miller, 1993; Simonetti et al., 2005; Jackson et al., 2004; Jeffries et al., 2003). Error propagation used follows that of Horstwood et al. (2003) where analytical errors of the unknown zircon analyses are combin ed with errors from the analyses of the standard FC-1 zircons by ( c / C)2 = ( mA / A)2 + ( B / B)2 (A-3) where A is the raw unknown zircon isotopic ratio and mA is the error of the mean of the analysis. B is the raw standard FC-1 zircon isotopic ratio and B is the standard deviation of all the standard FC-1 zircon analyses for a given analytical session. C is the drift and mass bias corrected unknown zircon ratio and c is propagated error for the unknown zircon analyses of interest. Error correlation factors were approximated between isotopic ratios simply by dividing their individual propagated errors. For example, an error correlation factor between the 206Pb / 238U and 207Pb / 235U error is obtained by dividing the 206Pb / 238U error by the 207Pb / 235U error (e.g., Horstwood et al., 2003). The drift and mass bias corrected isotopi c ratios and propagated errors were imported into Isoplot version 3.09a by Ludw ig (2004) where Tera-Wasserburg type and traditional concordia diagrams were constructed for zircon analyses of the two granitoid samples (WG04-114 and Ug-1) and the leuc osome sample WG05-02, respectively. 206Pb / 238U weighted mean age plots were also cons tructed for the granitoid samples. Note

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203 however, that isotopic data plotted on Te ra-Wasserburg concordia diagrams are not corrected for common lead. However, individual 207Pb / 235U and 206Pb / 238U ages were calculated from both common lead corrected and uncorrected data for comparison. For these data, the common lead correction wa s carried out using the method as outlined in Kosler and Sy lvester (2000), where common 206Pb is estimated using measured 207Pb / 206Pb ratios; the 207 method assumes c oncordance. For the leucosome sample, 207Pb / 235U and 206Pb / 238U ratios were corrected for common lead using measured 204Pb as described by Williams (1998). Note the common lead isotopic ratios used in both correction methods were approx imated using the twostage lead evolution model of Stacey and Kramers (1975). Furthermore, individual 207Pb / 206Pb ages were estimated from the uncorrected isotopic data of sample WG05-02 by interpolating between known 207Pb / 206Pb ratios and ages in Table 18. 3 of Faure (1986). A probability density plot was made for the 207Pb / 206Pb ages using Isoplot. Major and Trace Element Geochemistry Clean, representative portions of samp les DF02-114, WG04-114 and Ug-1 were first crushed to gravel-sized particles usi ng a Sturtevant rock Jaw Crusher. These fractions were then reduced to coarse sa nd using a Progressive Exploration Products mini Jaw Crusher model 150. The coar se sand fractions were then loaded into individual sealed ceramic cylinders along w ith two small ceramic powering balls. The sealed ceramic cases were then shaken in a Spex model 8000 mixer/mill for 30-60 minutes to create homogenized whole-rock powders from the samples. The whole-rock powders were subsequently sent to Acme Analytical Laboratories Ltd (Vancouver, B.C.) where both major and trace element analyses were carried out. In preparation for the analyses, the whole rock powders were first dissolved using a 4-acid

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204 digestion. Then major and trace element data were acquired using ICP-MS (inductively coupled plasma mass spectrometry). For de tails of the Acme ICP-MS analytical procedures see http://www.acmelab.com/cfm/index.cfm?It=100&Id=1

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205 APPENDIX B ACC LOWER PLATE LITHOLOGIC UNIT DESCRIPTIONS AND DISTRIBUTION General Statement Lower plate rock units mapped in the current study area fall into three distinct age categories: 1) Metasedimentary strata correlated with the Mesoproterozoic Belt Supergroup and Middle Cambrian formati ons, 2) Late Cretaceous intermediatecomposition intrusive rocks and 3) Middle Eocene intermediate-to-felsic-composition intrusive rocks. Notably, all Belt Supergroup and Middle Cambrian correlated metasedimentary strata in the study area have been subjected to middle to upperamphibolite grade metamorphism. In some areas, pronounced ductile deformation and tectonic attenuation of the metasedimentary strata accompanied this metamorphism (e.g., within the Lake of the Isle shear zone). In addition, in the easternmost part of the study area intrusive rock units were deformed under greenschist facies conditions (e.g., mylonitic granitoids of the Anaconda mylonite ). For simplicity, the details of such metamorphism and deformation are described and discussed in late r sections. Here, general lithologic descriptions, distinguishing characteristics and ove rall distribution of the mapped rock units are given to provide a metamorphic-stratigraphic framework for Belt Supergroup and Middle Cambrian corre lated metasedimentary strata and basic descriptions for the Late Cretaceous to Eocene intrusive rocks. Belt Supergroup and Middle Cambrian Met asedimentary Strata Correlation All metasedimentary strata exposed in the study area were originally mapped, described and correlated with the Belt Supe rgroup (then, the Belt Series) and Middle

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206 Cambrian section by Emmons and Calkins (1913 ). Much later, mapping and correlations made by Lonn et al. (2003) largely agree w ith the early work, a lthough Lonn et al. (2003) use more current Belt nomenclature for th is region (see Winston and Link, 1992 for a review and history of the Belt nomenclature). However, one di fference is that Lonn et al. (2003) group the Middle Cambrian formati ons into one undivided group, unlike Emmons and Calkins (1913) who mapped and correlated with individual formations. In this study, Belt correlation follows that of Lonn et al (2003), while Middle Cambrian-equivalent metasedimentary strata are ma pped and correlated with indi vidual formations as Emmons and Calkins (1913). Metamorphosed Belt Supe rgroup units mapped include (from old to young) the Greyson Formation, Ravalli Group, Helena Formation, and Missoula Group. Middle Cambrian-equivalent metasedimentary strata mapped include (from old to young) the Flathead, Silver Hill, and Hasmark Formations. Belt Supergroup-equivalent Metasedimentary Strata Pelitic Schist and Migmatitic Paragneiss The structurally lowest metasedimentar y rock unit exposed in the study area is largely comprised of a distinct reddish-bro wn to dark-brown or dark-grey schist correlated with the lower-Belt Greyson Formation. Because of its dark color, this schist unit is easily distinguished in the field, especi ally where in contact with relatively lightercolored intrusive rocks. In general, the schi st is fairly massive (and highly schistose), medium-grained containing abundant biotite wi th less (in decreasing abundance) garnet + sillimanite + quartz + plagioclase + K-feldspar muscovite. Excellent exposures of this rock unit are found along the walls of Sulliv an Creek and in the uppermost Tenmile Creek drainage. Exposures in the Tenm ile Creek headwall are characterized by alternating garnet-rich (smaller garnets 14 mm in diameter) and quartzofeldspathic

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207 horizons. Individual garnet-ric h horizons range in thickness from a few millimeters to a several centimeters. In the central mapping area, north of the c ontinental divide the schist grades to cordierite-bearing paragneiss. Good exposures of the paragneiss are found in several glacially polished outcrops s outhwest of the Lake of the Isle and in the uppermost reached of East Fork of the Twin Lakes Creek drainage. At these localities, the paragneiss is characterized by abundant sill imanite giving some outcrops a distinct battleship-grey hue. Mineral phases subordi nate to sillimanite include (in decreasing abundance) biotite + garnet + K-feldspar + quartz + plagio clase + cordierite. Note several outcrops of the paragneiss are migmatitic and show evidence for partial anatexis (melting) during severe ductile deformati on. These outcrops are characterized by boudinaged quartzite-rich horizons mesoscopic (nearly isoclinal) folds and transposed planar foliations; where granitic leucosome is found in boudin neck s or as concordant, thin elongate pods within the transposed fabrics Biotite Quartzite Paragneiss A prominent mica-rich quartzite unit overl ies the Greyson-correlated schist and paragneiss and is correlated with the Rava lli Group strata. The contact between these two units is observed only along the continen tal divide on the west ern flank of Mount Howe, where pelitic horizons of the Greyson gradually grade to more and more quartzite of the Ravalli. In most places, the meta-Rav alli Group strata consis t of medium-grey to dark-grey or dark-blue medium-grained qua rtzite paragneisses. These quartzite paragneiss are comprised of more or less pure qu artzite horizons that alternate with more pelitic (or schistose) horizons consisting of variable amounts of biotite (commonly

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208 chloritized) + K-feldspar + quartz + plagioclase sillim anite. Individual horizons (gneissic bands) range in thickness from a few millimeters to several meters in some places. These quartzite gneisses are also characterized by co mmon mesoscopic-scale folds, small-scale to outcrop-scale shear bands (mm to 10s of m thick) and rare boudins. Interestingly, some Ravalli outcrops preserve less deformed horizons with relic crossbedding between highly deformed horizons ; suggesting strain was distributed heterogeneously at some levels. The upper ~100-200 m section of the metamo rphosed Ravalli Group is distinctively more pelitic than lower sections. The be st exposure of the pe litic section is found adjacent to a small unnamed lake southeast of Storm Lake in the uppermost Twin Lakes Creek drainage. This pelitic horizon was also documented by both Emmons and Calkins (1913) and Lonn et al. (2003) ; the latter suggest correlation with the Saint Regis Formation of the uppermost Ravalli Group. Notably, Lonn et al. (2003) document kyanite in this upper pelitic ho rizon. Limited fresh kyanite was found at this locality but only seen in thin section (s ee description of the LISZ above). However, pseudomorphs resembling both kyanite and garn et are abundant in the more pelitic horizons of the uppermost Ravalli Group at this locality. Calc-silicate Paragneiss Above the Ravalli Group lies a thick seque nce of deformed, largely calc-silicate paragneiss correlated with the Helena Form ation (eastern equivalent to the Wallace Formation, both part of the Middle Belt Car bonate; Winston and Link, 1992). The calcsilicate paragneiss is typically a dense, fine-g rained and strongly layered rock. Light-tan to light-grey more silicic horiz ons comprised of calcite + quart z + chlorite alternate with

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209 pale-green to greenish-grey horizons mostly comprised of diopside + calcite (sometimes dolomitized) + tremolite (often highly altered) actinolite + quartz chlorite. These alternating horizons commonly range in thickness from a few millimeters to a few centimeters. In addition, some minor quartzite and biotite-ri ch schist horizons (usually ~1-5 cm thick) are found near the base of the metamorphosed Helena Formation, where the quartzite horizons are commonly boudi naged. Excellent exposure of the metamorphosed Helena Formation are found alo ng the pack trail on th e eastern flank of Mount Tiny and along the prominent ridge line west of Mount Haggin Notably, the Mount Tiny exposures are characterized by a di stinct cleavage that cuts the metamorphic foliation at a moderate angle (usually 35). This cleavage-foliation relationship geometrically resembles that of sedimentar y cross-bedding. Also note mesoscopic-scale folds are very common in the metamorphos ed Helena Formation exposures along the eastern flank of Mount Tiny, some exhibiting sheath-fold geometries. Calc-silicate Paragneiss and Biotite Schist Metasedimentary strata above the metamo rphosed Helena Formation and below the Middle Cambrian section are correlated with th e uppermost Belt Missoula Group. The Missoula group is typically subdivided into seve ral formations in other parts of western Montana (e.g., Snowslip, Shepard, Mt. Shie lds, Bonner, McNamara, Garnet Range and Pilcher Formations). However, in the current study area not all formations are expected to be present due to the progressively downcutting Cambrian unconformity in this region (Winston and Link, 1993 and references cited therein). In addition, because the Missoula Group strata present have undergone high grad e metamorphism and intense deformation this division is impractical and virtually impossible.

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210 Metasedimentary strata mapped as the Missoula Group are mostly comprised of calc-silicate paragneiss similar, in many aspe cts, to the underlying calc-silicate gneiss of the metamorphosed Helena Formation. Howe ver, the Missoula Gr oup calc-silicate gneiss is generally more diopside-rich, commonly cont ains chlorite and t ypically outcrops in a distinctively darker green hue. In addition, the calc-silicate gnei ss of the Missoula Group contains more interlayered schistose horiz ons than calc-silicate gneiss of the Helena Formation. The Missoula Group calc-silicate gneiss is well exposed along northwestern shore of Storm Lake where it is comprises the eastern limb of an antiformal structure Here, the calc-silicate gneiss exhibits prof ound ductile deformation and is characterized by extensive chocolate tablet-style boudinage of the more silicic horizons Note some metamorphosed Missoula Group strata are comp rised of schist and micaeous quartzite horizons, particularly near the bottom of the section. These horizons are usually characterized by alternating bi otite schist and biotite-rich quartzite horizons, where individual horizons usually range from 1 mm up to 50 cm in thickness. Some outcrops expose distinct blocky-sha ped quartzite boudins. Middle Cambrian-equivalent Metasedimentary Strata Micaeous Quartzite, Locally Mylonitic A thin quartzite unit overlies the metasedimentary strata correlated with the Missoula Group and is correlated with the Mi ddle Cambrian Flathead Formation. Outside the study area, the Flathead Formation is comprised of clean quartzite sandstone and interpreted to represent the first ma jor flooding surface to encroach on western Montana during the middle Cambrian transgre ssion (Winston and Link, 1993). However, within the study area, the Flathead Formation is highly metamorphosed and recrystallized to a coarse-grained, vitreous, and white to light-pink or tan pu re quartzite. Major

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211 outcrops of the metamorphosed Flathead Form ation are found directly west of Storm Lake on the western limb of an antiformal structure and on the eastern walls of the Sullivan Creek drainage in the southern study area. Notably, the Sullivan Creek exposure is strongly mylonitic and characterized by elongate quartz ribbons (up to ~ 1.5 cm in length) and abundant muscovite-fish, bot h comprising an east -southeast trending lineation. Lonn et al. (2003) and ONeill et al. (2004) interpret the Sullivan Creek quartzite to be a detached slab of meta morphosed and deformed Flathead Formation brought down against the Greyson schist (i.e ., a younger-on-older normal detachment). A similar structural relationship was docum ented at Short Peak, northeast of Sullivan Creek along the continental divide. Here, two smaller detached slabs of mylonitic quartzite correlated with the Flathead Formation rests on biotite granodiorite. Pelitic Garnet-biotite Schist In many places, a thin tan to light-brown colored schist unit overlies the Flathead Formation correlated quartzite and is correlated with the Middle Cambrian Silver Hill Formation. The Silver Hill schist is ge nerally fine to medium-grained containing abundant biotite defining a moderately define d foliation with lesser K-feldspar + quartz + plagioclase sillimanite. Small garnets ( 1-3 mm in diameter) are also common in some horizons. In addition, the uppermost horizons of the Silver Hill schist are calcareous and effervesce in HCL. Good exposur es of the Silver Hill schist can be found in the uppermost East Fork of the Twin Lakes Creek drainage and directly west of Storm Lake. Note at the Twin Lakes Creek locality the Silver Hill schist and other Middle Cambrian equivalent metasedimentary strata are folded into a tight, upright isoclinal antiform.

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212 White Marble, Locally Dolomitic The structurally highest metasedimentary unit mapped in the study area consists of a thin, course-grained white to light-grey locally dolomitic marble unit correlated with the Middle Cambrian Hasmark Formation. In several locations, this Hasmark marble can be found directly overlying th e Flathead and Silver Hill correlated metasedimentary strata, comprising a complete, yet metamorphos ed and highly deformed Cambrian marine transgressive sequence. Good exposures of the sequence are found on the southern flank of Mount Haggin within a larg e overturned nappe and on the wa lls of the Fourmile Basin Lakes drainage. The Hasmark marble is al so common as thin, but laterally extensive septa within the Storm Lake Stock. These marble septa are typically 10 m across but often extend for 100s to 1000s m through the intrusion. The uppermost Fourmile Basin Lakes drainage provides the most spectacular exposures of the marble septa. Here, several septa together define a west-northwest verging fold intruded by quartz diorite of the Storm Lake Stock Late Cretaceous Intrusions The Storm Lake Stock The Storm Lake Stock (SLS) is an interm ediate composite-type pluton that mostly outcrops over a ~ 5 km2 area within the western LISZ in the northeastern AnacondaPintlar Range, adjacent to Storm and Twin Lakes. The pluton was originally mapped by Emmons and Calkins (1913) who describe the SLS as undeformed medium-to-finegrained basic (mafic) granodiorite and quartzbearing diorite. Later, Wallace et al. (1992) described the pluton as being composed of several different phases including: quartz diorite, tonalite, granodi orite, quartz monzodiorite, and diorite. They refer to the pluton as the Stock of Storm La ke because of its size (or volume) and composite nature.

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213 Although some compositional variations were observed as described by Wallace et al. (1992), these variations are generally on the sm all scale. Therefore, here the SLS is subdivided into only two main phases: a granod iorite and quartz diorite phase. This subdivision (similar to the division made by Emmons and Calkins, 1913) was made because these two phases are by far the mo st voluminous and their distinction is important to later discussions involving the development of LISZ. The SLS granodiorite is was emplaced later than th e quartz diorite phase. The ag e relationship was observed in the uppermost Fourmile Basin Lakes drainage where the granodiorite phase contains several stopped blocks of the quartz dior ite phase. Documentation of this age relationship is important to later discussions concerning the evolution of the LISZ. The granodiorite phase of the SLS was pr eviously dated by Wallace et al. (1992) using the K-Ar method who report hornblende and biotite cooling ages of 116.4 4.6 and 78.7 1.6 Ma, respectively. They attribut e the anomalously old hornblende cooling age to extraneous argon and suggest a late Cret aceous age for the SLS granodiorite based on the biotite cooling age. In this study, two samples were collected from the SLS granodiorite for 40Ar/39Ar analyses to provide better age constrains. In addition, U-Pb zircon analyses were carried out on one of th e samples to provide a crystallization age for the SLS granodiorite phase. The SLS quartz diorite phase was sampled for 40Ar/39Ar and U-Pb zircon analyses. The SLS quartz diorit e phase had not been dated prior to this study. Storm Lake Stock Granodiorite The SLS granodiorite partly surrounds St orm Lake and lies west of the quartz diorite phase within the west ern LISZ. Several scattered outcrops, mostly along the

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214 eastern shores of Storm Lake, provide good exposur es of this phase. In this area, the SLS granodiorite is an undeformed medium-grain ed, salt-and-pepper, grey-colored rock comprised of the major mineral phases (in decreasing abundance) pl agioclase, quartz, biotite, hornblende and K-feld spar. In thin section, t ypical SLS granodiorite (sample WG04-114) is characterized by abundant subhe dral to euhedral pl agioclase phenocrysts ranging up to ~ 3-4 mm in length, often with well-developed albite twinning. Most of these plagioclase phenocrysts have a peppered appearance due to abundant fine-grained mineral inclusions. Biotite and hornblende phenocrysts commonly range up to ~ 2-3 mm in length and some larger hornblendes are twin ed. Notably, most biotite and hornblende phenocryst are anhedral in form, having irregular or ra tty grain boundaries. K-feldspar and quartz are generally pres ent as relatively small ( 1.5 mm) phenocrysts. Accessory mineral phases identified in thin section include (in decreasing abundance) magnetite, apatite, zircon and titanite. These phases are found as small inclusions within the phenocrysts phases. Several good exposures express the relati onship between the SLS granodiorite and highly deformed Belt Supergroup/middle Cambri an-correlated metasedimentary strata in the western LISZ. Immediately west-northwest of the Storm Lake dam (north end of the lake) the undeformed granodiorite phase obl iquely crosscuts deformed calc-silicate gneiss, quartzite and schist on th e northern tip of an antiformal structure. These strata are correlated with the Missoula Gr oup, Flathead and Silver Hi ll Formations, respectively (see descriptions above). A similar relations hip can be observed at the divide between Storm and Twin Lakes Creek drainages, ~2 km southwest of Storm Lake. At this locality, the granodiorite cu ts orthogonally across a well-developed metamorphic

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215 foliation in deformed biotite quartzite and calc-silicate gneiss (Ravalli Group and Helena Formation equivalents, respectively) on th e northeastern limb of the same antiform structure. In addition, th e SLS granodiorite contains a bundant thin, elongate septa of deformed white marble correlated with the middle Cambrian Hasmark Formation. These septa are exposed on the eastern shores of Storm Lake and to the southeast along the northeastern wall of th e upper Storm Lake drainage. A few of the most extensive Hasmark septa within the SLS granodiorite we re mapped and are shown in Storm Lake Stock Quartz diorite A quartz diorite phase comprises most of the eastern SLS and is mainly exposed over a broad area in the upper Fourmile Ba sin Lakes and upper Twin Lakes Creek drainages. Here, the SLS quartz diorite is an undeformed medium-to-coarse-grained grey-colored rock somewhat similar looking in outcrop to the SLS granodiorite (i.e., grey, salt-and-pepper appearance). However, the quartz diorite is typically more coarsegrained and contains more plagioclase a nd hornblende; plagiocl ase and hornblende comprise over 70 % of the rock at many exposures. Other major mineral phases found in the quartz diorite include (in decreasing abunda nce) biotite, quartz and K-feldspar. A thin section of the SLS quartz diorite (sample WG04-052 from upper Fourmile Basin Lakes drainage) exhibits abunda nt large euhedral plagiocl ase phenocrysts ranging in size up to ~ 4-5 mm in longest dimension. Most plagioclase phenocry sts in thin section display a prominent albite twinning. Hornble nde is present as large, sometimes elongate phenocrysts up to ~ 5-7 mm in length. Bi otite is observed of ten as lath-shaped phenocrysts ranging up to 3-4 mm in length. No tably, the hornblende and biotite of the quartz diorite are subhedral-euhedral in cr ystal form, therefore having a more defined

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216 form than hornblende and biotite of the SL S granodiorite. K-fe ldspar and quartz are present as relatively small phenocrysts up to ~ 2-3 mm in diameter and some K-feldspar phenocrysts exhibit tartan twinning. In addi tion, myrmekitic worm-like intergrowths of quartz and K-feldspar are commonly found on pl agioclase phenocryst rims adjacent to Kfeldspar. Accessory mineral phases identified in the SLS quartz diorite and their relative abundances are similar to those of the SL S granodiorite phase (i.e., in decreasing abundance: magnetite, apatite, zircon and titanite). Exposures along the western walls of the Fourmile Lakes Basin drainage, just above several small lakes express the relati onship between the SLS quartz diorite and the LISZ in the western study ar ea. At this locality, the undeformed SLS quartz diorite clearly crosscuts highly defo rmed metasedimentary strata correlated with the upper Belt Missoula Group and the middle Cambrian se ction. Emmons and Calkins (1913) document this same crosscutting relationship. In addition, in the uppermost Fourmile Basin drainage along the eastnortheast facing headwall, se veral Hasmark marble septa together outline a northwest verging fold. Emplacement of the undeformed SLS quartz diorite clearly post-dates folding of the septa. The Deformed Quartz Diorite Sill East of the widespread exposures of quart z diorite within upper Fourmile Basin and Twin Lakes drainages, a deformed narrow ( 1 km thick) quartz diorite sill was mapped The quartz diorite sill is emplaced within steeply dipping, strongly metamorphosed and deformed Belt Supergroup/middle Cambrian-cor related metasedimentary strata in the central and eastern LISZ and can be trace eastwa rd to the base of Mount Haggin. Unlike the exposures of undeformed quartz diorite in the west, the quartz diorite sill exhibits a well-developed solid-state foliation concordant with the metamorphic foliation within the

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217 adjacent Belt and middle Cambrian-equivalent metasedimentary country rocks in the LISZ. The foliation in the quartz diorite sill is defined by the pa rallel alignment of elongate hornblende clots, flattened plagiocl ase and K-feldspar phenocrysts, and quartz ribbons (all 4-5 cm in length). Some larger (~.5 to 1 cm diameter) plagioclase phenocrysts exhibit symmetric and asymmetric sigmoidal geometries as well. The foliated SLS quartz diorite sill is well-exposed along the southern side of the upper East Fork of the Twin Lakes drainage (east-northeas t of the Lake of the Isle) and in the upper reaches of the Mill Creek drainage. At the upper East Fork of the Twin Lakes drainage exposure, the sill contains thin ( 5 m thick) septa of biotite-rich quartzite gneiss and calc-silicate gneiss; the latter resembles calc -silicate gneiss of me tamorphosed Missoula Group. In the eastern study area, a fairly large isolated pluton also comprised of quartz diorite is exposed on Mount Haggin. Th e quartz diorite on M ount Haggin crosscuts highly deformed and metamorphosed metasedi mentary strata correla ted with the upper Belt Supergroup and middle Cambrian sect ion along the upper limb of a large, eastverging recumbent nappe structure (Heise, 1983) Therefore, the qua rtz diorite intrusion on Mount Haggin shows the same relationshi p with the deformed and metamorphosed metasedimentary wall rocks as the quartz dior ite of Fourmile Basin Lake drainage in the west. However, the quartz diorite on Mount Haggin is deformed and overprinted by a mylonitic fabric in some places, presumably the same fabric that overprints the granitoids in the mylonite zone to the east (Heise, per comm.). Eocene Intrusions The central, eastern, and southern parts of the study area, in the northeastern Anaconda-Pintlar Range, are characterized by highly voluminous Eocene intrusive rocks.

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218 In this region, four major intrusive phases were identified and mappe d as separate units. These intrusive rock units include: medium -grained biotite granite, medium-to-coarsegrained two-mica (biotite and muscovite) gran ite, biotite granodiorite, and porphyritic two-mica granite. A number of northeast-tre nding dacite dikes were also mapped in the central part of the study area Crosscutting relationships observed in the field indicate the following order of emplacement for the mapped intrusive units, from oldest to youngest: biotite granite two-mica granite > biotite gr anodiorite > porphyritic twomica granite dacite dikes. These subdivisions a nd relative ages agree with the early work of Emmons and Calkins (1913). Howeve r, they also describe and map a porphyritic biotite granite phase, placing this unit betw een the biotite granodi orite and porphyritic two-mica granite in relative age (see their manuscript for a very complete assessment of the intrusive rocks of the northeastern Anaconda -Pintlar Range). Here, relatively basic descriptions are provided fo r the different intrusive units mapped in the current study area; emphasis is placed on distinctive charac teristics and general spatial distribution of the rock units throughout the st udy area. Note the biotite and two-mica granite in the easternmost study area have greenschist my lonite fabrics (i.e., within the Anaconda mylonite). The characteristics of the greenschist mylonite fabrics are described in a later section. Biotite Granite Biotite-bearing granite outcrops over a larg e area within the s outheastern portion of the study area, reaching from the Tenmile Creek drainage to the upper Clear Creek drainage within the Anaconda mylonite. Wh en not mylonitized, the biotite granite is typically a fabric-less medi um-grained rock largely comprised of (in decreasing

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219 abundance) quartz + K-feldspar + plagioclase + biotite. Quartz typi cally forms anhedral to subhedral phenocrysts ranging in diameter up to ~3-4 mm. Plagioclase is similar in size to quartz, but usually subhedral to euhe dral. Plagioclase comm only exhibits albite twinning in thin section. In addition, K-feldspar exhibit ta rtan twinning in thin section and some phenocrysts range up to ~8 mm in longest dimension. Individual lath-shaped biotites are usually ~1-3 mm in length and of ten found in randomly oriented clusters with other biotite laths. At high elevations (e.g., along the northw est-trending ridge th at separates the Tenmile and Mill Creek drainages) the biotite granite is highly chloritized, but lacks a well-developed mylonitic fabric. At these high elevation outcrops the biotite granite also exhibits a pervasive green hue and is brit tlely fractured; presumably because high elevations approach the previous detachment level. On the eastern side of the Mill Creek drainage, within the mylonitic biotite granite, a slab of coarse white-pink quartzite is juxtaposed with the biotite granite along a lowangle detachment fault. This quartzite has been correlated with the middle Cambrian Flathead Formation by Emmons and Calkins (1913). However, without any adjacent metasedi mentary units present it is difficult to confidently make such a correlation. Two-mica Granite Medium-grained two-mica granite is the most voluminous intrusive rock type mapped within the study area. This rock type outcrops in several places: north and east of Mount Haggin, including the upper Cl ear Creek drainage, the upper Mill Creek drainage, a large glacial cerk south of the Lake of the Isle, and along the northwesttrending ridge that separates the Sullivan and Seymour Lakes Creek drainages in the

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220 southwestern study area. The exposures nor th and east of Mount Haggin are mostly mylonitic. The non-mylonitic two-mica granite is remarkably similar to the biotite granite in both texture and mineralogy; the only difference is the presence of muscovite and sometimes slightly more abundant quartz a nd K-feldspar in the two-mica granite. In thin section, muscovite is found as randomly oriented, some times clustered, lath-shaped crystals similar in size to usually more abunda nt biotite laths (~1-3 mm laths). Biotite is sometimes chloritic. Notably, in a thin se ction of one two-mica granite sample (WG04043), quartz and K-feldspar show substantial de formation, which could not be detected at hand sample or outcrop scale. In this thin section, quartz exhibits strongly developed undulatory extinction and is sometimes character ized by multiple subgrains with serrated grain boundaries within (i.e., grain boundary migr ation). In addition, in K-feldspar is highly fractured sometimes exhibiting a comple x, anastomosing pattern of fine-grained factures. Biotite Granodiorite A medium-grained biotite granodiorite co mprises a west-dipping tabular sheet-like intrusion that stretches from the upper Tenm ile Creek drainage, acr oss Mill Creek into the upper Clear Creek drainage. In addition, a large pluton in the Seymour Lakes Creek drainage in the southwestern part of the study area is comprised of a similar biotite granodiorite (Lonn et al., 2003). In th e upper Clear Creek drainage the biotite granodiorite narrows to a fi nger-like extrusion that extends into the two-mica granite and is strongly mylonitic. However, elsewher e the biotite granodiorite is observed to be non-mylonitic and mostly comprised of (in decreasing abundance) quartz + plagioclase + K-feldspar + biotite hornblende (rare).

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221 Similar to the biotite granite, the biotite granodiorite is observed to be brittlely fractures and mildly to modera tely chloritized at some hi gh elevations exposures. In particular, along the northwest-trending ridge which separates the Tenmile and Mill Creek drainages (southwest of Short Peak) the biotite granodiori te is cut by several closely spaced (~5-30 m apart) high angle br ittle faults. Individual fault blocks are moderately chloritized and th e degree of chloritization generally increases towards the faults within individu al fault-bound blocks. Porphyritic Two-mica Granite A porphyritic two-mica granite pluton wa s mapped in the upper Twelvemile and Tenmile Creek drainages. This intrusion wa s originally mapped and referred to as the Twelvemile Batholith by Emmons and Calk ins (1913). Mineral ogically, the porphyritic two-mica granite is very similar to the to medium-grained two-mica granite and contains (in decreasing abundance) quartz + K-feldspar + plagioclase + biotite + muscovite. Kfeldspar is commonly more abunda nt than plagioclase feldspar and biotite is usually more abundant that muscovite. In addition, some exposures exhibit very large tabular Kfeldspar phenocrysts ranging up to ~3-4 cm in longest dimension and quartz phenocrysts up to ~1-1.5 cm in diameter; the larger K-feldspar and quartz phenocrysts are often concentrated in clusters and surrounded by relatively finer-graine d quartz, K-feldspar, plagioclase, biotite and muscovite. In th e upper Tenmile Creek dr ainage several thin (typically 1-5 m) aplite dikes extend from th e main body of the porphyritic two-mica granite pluton. These aplite dikes are sole ly comprised of (in decreasing abundance) Kfeldspar + quartz + muscovite. Muscovite is very coarse-grained in within these dikes

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222 (sometimes up to ~1-1.5 cm flakes). Some aplite dikes trend to the north while others trend east or northeast. Northeast-trending Dacite Dikes Numerous thin (~1-5 m thick) medium to dark-grey dacite dikes cut through the central part of study area and predominantly tr end to the northeast. These dacite dikes are characterized by aphanitic porphyritic te xture. In thin section, plagioclase phenocrysts are usually subhedral to euhedral in crystal fo rm and ~2-10 mm in diameter; many exhibit albite twinning and concentric zoning patterns. Some plagioclase phenocrysts are highly angular and apparent fragments of original whole plagioclase phenocrysts. Quartz is present as less abundant sub-rounded to rounded phenocrysts usually ~1-2 mm in diameter. Markedly, some quartz phenocrysts exhibit welldeveloped undulatory extincti on and internal subgrains with serrated grain boundaries. Biotite is common as small (~1 mm by 2 mm ) lath-shaped phenocrysts dispersed among the larger plagioclase and quartz phenocrysts; some biotite laths are moderately chloritic. These phenocrysts phases are separated by a ve ry fine-grained matrix comprised mostly of plagioclase with lesser quartz and biotite. Although not shown on the map in Appendi x E, these dikes are especially concentrated within the metamorphosed Greyso n Formation (pelitic schist and migmatitic paragneiss) in the area north of Mount Evans an d south of the Lake of the Isle (see the mapping of Emmons and Calkins, 1913 and Lonn et al., 2003). Most of the dacite dikes mapped in this study trend to the northeast and cut across the Belt and middle Cambrianequivalent metasedimentary strata in the LISZ at high angle. Just to the northeast of the Lake of the Isle the dacite dikes crosscuts the deformed quartz diorite sill, the

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223 metamorphosed middle Cambrian section, and calc-silicate an d biotite quartzite paragneisses correlated with the metamorphos ed Helena Formation and Ravalli Group, respectively. One mapped dacite dike crossc uts the western part of a two-mica granite pluton south of the Lake of the Isle Note the dacite dikes mapped in this study area are likely the northern continuati on of the extensive NE-trending middle Eocene dike swarm in the southwestern Anaconda-Pintlar Range and the Chief Joseph Batholith documented by previous workers (e.g., Desmarais, 1983; Hyndman et al., 1988; Wallace et al., 1992).

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224 APPENDIX C DESCRIPTION OF SAMPLE ME-231 ME-231 was chosen for the thermobarometry of this study. This sample was taken from one of several large outcrops exposing highly deformed migmatitic pelitic paragneiss (metamorphosed Greyson Fm., Lower Be lt) near the Lake of the Isle, in the central part of the LISZ (Appendix F). ME -231 bears an assemblage of (in increasing relative abundance) cordierite + garnet + K-feldspar + albite + biotite + quartz + sillimanite; this upper-amphibolite facies asse mblage is ideal for calculation of several mineral phase equilibria used in conventional thermobarometry in pelitic rocks (e.g., Spear, 1993). The outcrop from which ME -231 was taken also contains abundant granitic leucosome commonly f ound in boudin neck pressure sh adows; this is consistent with the interpretation of insitu partial anatex is occurred during ductile deformation in the central LISZ. Furthermore, the migmatitic pe litic paragneiss here are the most intensely deformed and highly metamorphosed exposed in the LISZ. Therefore, PT estimates calculated from the equilibrium phase assembla ge of ME-231 most likely represent peak pressure and temperature condi tions (or closest to peak PT possible) associated with upper-amphibolite facies metamorphism and ductile deformation of Belt and middle Cambrian-equivalent metasedimentar y strata in the LISZ. More specifically, the outcrop from whic h ME-231 was collected is characterized by highly transposed, planar metamorphic fabr ics. Abundant sillim anite mats and lesser biotite are conspicuous and define a planar sub-vertical to vertical foliation in outcrop (and many other outcrops in the central LISZ). Quartzite-rich horizons alternate with

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225 pelitic horizons. The quartzite -rich horizons are moderately to highly flattened and often boudinaged to thin ( 10 cm thick) elongate blocky to el lipsoid boudins. The longer axes of these boudins are usually oriented parallel to the steep transposed and planar foliation in outcrop. These quartzite-r ich horizons are also found folded into elongate mesoscopicscale isoclinal, sometimes rootless folds with hi nge lines oriented parallel to the planar fabrics. In addition, light-grey to tan or lig ht-orange colored albite + quartz + k-feldspar bearing leucosome is common in boudin necks (i.e., between individu al boudins) and as thin ( 5 cm) elongate pods ( 0.5 m) within the planar fabrics of the outcrop. In thin section, sample ME-231 exhib its a pervasive metamorphic foliation comprised of aligned fresh sillimanite fi brolite mats on a less abundant light-brown biotite substrate (i.e., epitaxal sillimanite). The individual sillimanite fibrolite and biotite layers of this foliation are commonly 2 mm thick and generally spaced by ~ 4 to 8 mm. Thin, elongate ( 1.5 mm long) illmenite ribbons are commonly associated with sillimanite fibrolite mats, with their long ax es parallel to fibrolite mats and the overall orientation of the foliation in ME-231 (see photomicrographs of ME-231 in the description of LISZ above). The sillimanite-biotite foliation of ME-231 encases a population of garnet, cordierite and K-feldspar porphyroblasts. All of these porphyrobl asts show a pretectonic relationship to the foliation where the foliation wraps (or bent) around individual porphyroblasts (Yardl ey, 1999, p. 171). Because of this relationship, pressure shadows filled with fine-grained quartz, alb ite and K-feldspar are commonly associated with the larger porphyroblasts (mostly ga rnet and cordierite) in ME-231.

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226 The garnet porphyroblasts of ME-231 ar e typically anhedral to subhedraldodecahedrons and mostly range from ~1 to 10 mm in diameter. A few larger garnets are tabular-shaped with their longer axes oriented parallel to the sillimanite-biotite foliation. Note all garnet porphyroblasts are slightly to moderately fractured, where fractures are oriented near 90 to the encasing sillimanite-bio tite foliation. In addition, some garnets exhibit minor embayments along grain boundaries. Cordierite in ME-231 is common as la rge, slightly flattened sigma-type porphyroblasts ranging from ~ 4 to 12 mm in diameter. These sigmoidal-shaped cordierites are symmetric or slightly asymmetric; many havi ng tails sub-parallel to the sillimanite-biotite foliation described above. Kfeldspar is found as moderately to highly flattened, elongate asymmetric porphyroblasts or ribbons with larg e aspect ratios (e.g., ~1-2 mm long by 0.2 mm thick). Some of th e most severely flattened K-feldspar porphyroblasts exhibit subgrains and moderate undulatory extinc tion. In addition, some K-feldspar porphyroblasts in ME-231 have e xhibit areas of myrmekitic intergrowths with quartz. Both garnet and cordierite porphyroblasts in ME-231 contain inclusions suites. Garnets have abundant quartz and rare biotite inclusions, both usually up to ~ 1 mm in diameter. Cordierite porphyroblasts commonly contain several small ( .2 mm in diameter) zircon inclusions marked by bright yellow damage halos. A few cordierite porphyroblasts also host small bi otite inclusions typically .5 mm in length. Note both garnet and cordierite porphyroblas ts in ME-231 lack albite plag ioclase inclusions. Fine-grained hair-like sillimanite fibrolite inclusions are abundant in both garnet and cordierite porphyroblasts. In the garnets, these inclusions define a preserved internal

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227 sub-planar foliation oriented at high angles (often ~90) to the pervasive external sillimanite-biotite foliation. The smallest garnets (~1 mm diamet er) exhibit hair-like sillimanite fibrolite inclusions as well but the fibrolites are generally arranged in small, circular wreath-like patterns. The diameter of the hair-like sillimanite wreaths are commonly about half that of the host garnet. The hair-like silliman ite fibrolites also define a preserved internal foliation in co rdierite porphyroblasts, somewhat analogous to that of the larger ga rnets porphyroblasts. However, the in ternal foliation in cordierites is commonly wavier and more chaotic, somewh at analogous to the classic syn-tectonic snowball texture of garnet s described by Spry (1963). Albite and quartz are common in quartzof eldspathic layers (~1 to 8 mm thick) between the sillimanite fibrolite mat-biotite layering of the foliation described above. In the feldspathic layers individual albite and quartz grains are usually fine-grained and anhedral, often ~ 0.1 to 0.8 mm in diameter Both albite and quartz grains display undulatory extinction, this phenomenon being most intense in the later of the two. The deformed quartz grains often have several subgrains, some exhibiting serrated grain boundaries with respect to other quartz grains. In addition, quartz can be found as thin, highly flattened elongate quartz ribbons often reaching lengths of ~ 4 to 8 mm. More fresh and coarser grained albite, quartz and Kfeldspar are also found as thin (~1 to 1.5 mm) elongate (up to few cm in length) leucosome lenses or fingers between sillimanite-biotite and feldspathic layers of the foliation describe d above. Individual albite, quartz and K-feldspar grains are typical ly subhedral to euhedr al in shape, ranging in diameter from ~ 0.5 to 1. 5 mm. Albite and K-feldspar grains exhibit well-developed albite twinning and tartan twinning, respectiv ely. Furthermore, the individual albite,

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228 quartz and K-feldspar grains of the leucosome fingers do not exhibit undulatory extinction, subgrains or irregular grai ns boundaries as those described in the quartzofeldspathic layering. Notably, the placement of the albite + quartz + K-feldspar bearing leucosome viewed in thin sections of ME-231 is very similar to its macroscopic placement in outcrop. In thin section, the leucosome is commonly found in both garnet and cordierite porphoroblast pressure shadows, analogous to the leucosome found within boudin necks pressure shadows in outcrop. In addition, some garnets porphyroblasts display thin fingers, or veinlets, of the leucosome cross-cu tting the grains, parall el to the foliation of ME-231. Individual lenses of the leucos ome are also found within the sillimanite fibrolite mat-biotite layers as ~1 mm to ~1 cm in length (see photomicrographs in description of LISZ above). Muscovite is absent from ME-231 except for a few small, rare isolated clusters. These muscovite clusters are characterized by fresh muscovite laths, where individual laths reach ~ 1 mm in length. In all case s, the muscovite crosscuts the sillimanite fibrolite mat-biotite foliation of ME-231 at m oderate to high angles. Rare fine-grained muscovite is found along fractures in a few la rger garnet and cord ierite porphyroblasts. In addition, chlorite mica is absent from ME-231.

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229 APPENDIX D MAJOR AND TRACE ELEMENT GEOCHEMISTRY Purpose Major and trace element analyses were carried out to investigate the genetic relationship between late Cret aceous intrusions found within the ACC. Three samples were chosen for these analyses. DF02-114 wa s collected from a large medium-to-coarsegrained undeformed granodiorite block in th e detached ACC upper plate located ~3 km east of the greenschist facies mylonite zone exposed in the Mill and Clear Creek drainages. Notably, the upper plate granodiorite is similar to the undeformed granodiorite phase of the Storm Lake Stock (S LS) located in the western part of the ACC lower plate (Appendix F). In particular, the upper plate granodiorite sample DF02-114 is similar to WG04-114, a sample collected from the SLS granodiorite. These two granodiorite samples differ only slightly in texture and mineralogy, where DF02-114 is coarser-grained and contains less hornblende, le ss plagioclase, and more K-feldspar. In addition, these two samples gave similar biotite 40Ar/39Ar cooling ages; DF04-114 and WG04-114 gave biotite 40Ar/39Ar cooling ages of 76.3 1.1 Ma and 74.0 0.9 Ma, respectively (Chapter 6). The similar mineralogy, texture, and biot ite cooling ages of sample DF02-114 and WG04-114 indicate that the upper plate granodi orite and the SLS granodiorite are likely genetically related and originally part of th e same pluton. One possibility is the upper plate granodiorite represents a slightly more fractionated and evolved phase of SLS granodiorite emplaced at a shallower depth. This interpretation could account for the

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230 slightly more evolved nature of the upper plate granodiorite (i.e., as the upper plate granodiorite rose from the SLS granodiorit e below it fractionated and become more evolved). In addition, emplacement into th e shallower (and cooler) crust above the SLS granodiorite could also account for the sligh tly older biotite cooling age obtained for DF04-114 (i.e., the top of the pluton would have cooled sligh tly earlier). It is possible the upper plate granodiorite and the SLS granodiorite were originally part of the same pluton that was emplaced into the ACC lower plate in the late Cretaceous. Subsequently, the shallower part of the pluton (the upper plate granodiorite) was detached and displaced to the east-s outheast along the Anaconda detachment during the Eocene. If this relationship is proven plausible then the map distance between the upper plate granodiorite and SL S granodiorite can be used to estimate the amount of horizontal displacement facilitated by the Anaconda detachment since the onset of extension at ~53 Ma. Major and trace elem ent data were also obtained from Ug-1, a sample from the deformed quartz diorite si ll emplaced within th e attenuated (ductily thinned) Mesoproterozoic and middle Cambrian -equivalent metasedimentary strata in the central and eastern LISZ (Appendix F). These da ta are needed to test the quartz diorite sill as a possible parent for the upper plate gr anodiorite (i.e., the upper plate granodiorite could have originated from the quartz diorite sill by fractional crystallization rather than the SLS granodiorite). Major and Trace Element Analytical Results Whole rock powders were prepared from samples DF02-114, WG04-114 and Ug1 using traditional rock crushing and powderi ng methods and sent to Acme Analytical Laboratories Ltd for ICP-MS (inductively c oupled plasma mass spectrometry) major and trace element analyses. Methods for sample preparation are summarized in Appendix A.

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231 See http://www.acmelab.com/cfm/index.cfm?It=100&Id=1 for Acme ICP-MS analytical procedures. Major Element Analyses The results of the major elemental analyses from samples DF02-114, WG04-114 and Ug-1 are reported in Table D-1 and plotte d on Harker variation diagrams in Figure D-1. All major element oxides are reported in weight percent. The results show that all three samples are intermediate (55-62 wt % SiO2) in composition. The upper plate granodiorite DF02-114 is the most evolved, followed by the granodiorite WG04-114, and then the quartz diorite sill Ug-1 (SiO2 wt % = 64.0, 59.5 and 59.2, respectively). All three samples are high in K2O (2.27-3.74 wt %) as well fall ing in the high-K calc-alkaline field of the subalkalic rock classification. In additi on, DF02-114, WG04-114 and Ug-1 are very similar in major element compositi on (also see Table 1). Sample DF02-114 is somewhat similar to Ug-1 and WG04-114 in so me major elements, plotting to the right of WG04-114 and Ug-1 on Harker variation diag rams. However, DF02-114 is slightly depleted in Al2O3, Fe2O3, MgO, and CaO and more enriched in K2O with respect to Ug-1 and WG04-114. Trace Element Analyses The results from trace element analyses are summarized in Table 1 as well. These trace element concentrations are reported in ppm (parts-per-mill ion). Select trace elements concentrations from each sample were also plotted in Figure 2a, which displays an incompatible trace element spider diagram. On the spider diagram the trace elements were plotted from left to right in order of increasing mantle compatibility (e.g., see Rollinson, 1993, p. 144). Trace element concentrations for the average upper and lower crust are also plotted in Figur e D-2a for comparison to the un knowns of this study. The

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232 Table D-1. Results from major and trace element analyses Sample DF02-114WG04-114 Ug-1 Rock type hb bt granodioritehb bt granodiorite q. diorite Major Elements (wt %) SiO2 64.0159.52 59.24 Al2O3 15.9416.26 17.13 Fe2O3 4.957.13 6.76 MgO 2.173.25 3.13 CaO 4.415.92 6.15 Na2O 3.043.04 3.38 K2O 3.742.75 2.27 TiO2 0.610.83 0.76 P2O5 0.170.24 0.19 MnO 0.070.11 0.11 Cr2O3 0.000.00 0.00 LOI 0.700.70 0.60 TOT/C 0.060.04 0.03 TOT/S 0.010.01 0.01 Total 99.8199.76 99.73 Trace Elements (ppm) Ba 741.80787.00 917.90 Be 1.002.00 <1 Co 12.2020.50 19.40 Ni 11.0029.00 22.00 Sc 11.0016.00 15.00 Cs 6.804.20 1.80 Ga 17.0019.20 17.80 Hf 6.105.50 4.60 Nb 10.7010.80 8.10 Rb 133.00111.60 61.40 Sn 2.001.00 1.00 Sr 515.00636.60 470.40 Ta 1.000.70 0.70 Th 19.808.00 4.30 U 2.401.60 1.00 V 99.00147.00 108.00 W 12.006.20 9.60 Zr 203.60215.70 181.80 Y 21.6023.10 19.20 La 55.0028.00 26.60 Ce 112.5065.40 57.00 Pr 10.277.18 5.56 Nd 34.9029.20 21.80 Sm 5.805.30 4.30

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233 Table D-1. Continued. Sample DF02-114WG04-114 Ug-1 Rock type hb bt granodioritehb bt granodiorite q. diorite Trace Elements (ppm) Eu 1.181.39 1.27 Gd 3.774.61 3.53 Tb 0.670.64 0.57 Dy 3.423.81 3.24 Ho 0.780.78 0.67 Er 2.132.18 1.90 Tm 0.320.38 0.31 Yb 1.952.05 1.79 Lu 0.370.33 0.33 Mo 0.501.10 <.1 Cu 14.5027.30 13.60 Pb 4.101.60 1.50 Zn 22.0071.00 48.00 Ni 7.2011.80 8.00 As 0.800.80 0.70 Cd <.10.10 <.1 Sb <.1<.1 <.1 Bi <.1<.1 <.1 Ag <.1<.1 <.1 Au 0.501.00 1.00 Hg <.01<.01 0.01 Tl 0.300.40 0.30 Se <.5<.5 <.5 Major element ratios Al2O3 / (Na2O+K2O+CaO) 1.421.39 1.45 Trace elements ratios Zr/Y 9.439.34 9.47 Ce/Yb 57.6931.90 31.84 Note: q = quartz, hb = hornblende, bt = biotite average upper and lower crust trace element compositions were taken from Taylor and McLennan (1981) and Weaver and Tarney (1984), respectively. In addition, all trace element data plotted in Figure D-2a were normalized to the primitive mantle trace element concentrations of Mc Donough et al. (1991).

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234 Overall, samples DF02-114, WG04-1 14, and Ug-1 yielded trace element concentrations similar to the average upper crust. The trace element trends for the three samples lack pronounced negative rubidium, thorium, uranium, and yttrium anomalies typical of the depleted lowe r crust. DF02-114 yielded si gnificantly higher thorium, lanthanum, and cerium concentrations than the average upper cr ust and WG04-114 and Ug-1 gave thorium and uranium concentrations lower than the upper crust. In addition, the rubidium concentration of Ug-1 is lower than rubidium concentration in the average upper crust (Fig. 2a). Rare earth elements (REE) concentrations for samples DF02-114, WG04-114, and Ug-1 are presented in Figure 2b and normalized to the primitive mantle REE concentrations of McDonough et al. ( 1991). Overall, all three samples are characterized by moderately steep REE tr ends with negative slopes indicating an enrichment and depletion in th e light REE (left) and heavy REE (right), respectively (the Ce/Yb ratios for the three samples ranges from 57.7-31.8). Such patterns are indicative of magma source rocks that contain garnet and/or amphibole which strongly partition heavy REE relative to light REE (Rollinson, 1993). WG04-114 and Ug-1 yielded very similar REE patterns (Fig. 2b). The REE trend for WG04-114 closely mimics the REE trend of Ug-1, lying just above it indicating an overall enrich ment the all the REE. The DF02-114 REE pattern is similar to the R EE patterns of WG04-114 and Ug-1 over the medium to heavy REE (right to center, Fig. 2b). DF02-114 is more en riched in the light REE. In addition, DF02-114 exhibits a slig htly negative anomaly in europium suggesting equilibrium with fractionation of plagio clase (Best and Christiansen, 2001). The major and trace element data obt ained for DF02-114, WG04-114, and Ug-1 suggest these samples could be related by pro cess of fractional crysta llization (Rollinson,

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235 1993). In particular, the similar REE patte rns of WG04-114 and Ug-1 suggest the SLS granodiorite could have formed by fraction crystallization from the quartz diorite sill. Likewise, these data suggest the upper plate granodiorite (DF02-114) could have fractionated from sources similar the SLS granodi orite or the quartz dior ite sill given that the fractionating mineral phases acted enrich ed the daughter magma in light REE while leaving the heavy REE la rgely unchanged.

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236 Figure D-1. Major element Harker variati on diagrams for samples Ug-1, WG04-114, and DF02-114.

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237 Figure D-1. Continued.

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238 Figure D-1. Continued.

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239 Figure D-2. REE trend and trace element sp ider diagrams for samples Ug-1, WG04-114, and DF02-114. A) All three samples show enrichment in light REE relative to heavy REE. WG04-114 and Ug-1 show very similar trends in REE while DF02-114 is more enriched in La, Ce, Pr and Nd. B) All three sample show trace element concentrations largely similar to the average upper crust.

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240 APPENDIX E THERMOCALC AND AX OUTPUT FILES AX OUTPUT: Calculations for P = 4.0 kbar and T = 800C ________________________________________________________________________ g garnet 2-site mixing + Regular solution gammas Ferric from: Cation Sum = 8 for 12 oxygens W: py.alm=2.5, gr.py=33, py.andr =73, alm.andr=60, spss.andr=60 kJ oxide wt % cations activity sd % SiO2 36.88 2.934 py 0.0025 0.00111 45 TiO2 0.14 0.009 gr 0.000048 0.0000268 56 Al2O3 21.32 2.000 alm 0.47 0.070 15 Cr2O3 0.16 0.010 spss 0.0000014 0.00000080 59 Fe2O3 2.71 0.162 andr FeO 35.37 2.353 MnO 0.50 0.034 MgO 2.92 0.346 CaO 1.13 0.097 Na2O 0.30 0.046 K2O 0.11 0.011 Totals 101.56 8.000 _______________________________________________________________________ bi rim Al-M1 ordered, site-mixing model + macroscopic RS gammas: (ann, phl, east, obi) Ferric from: Tet + Oct cation sum = 6.9 for 11 oxygens. Max Ratio = 0.15 SF model parameters: Wpa=9, Wpe=10, Wpo=3, Wao=6, Wae=-1, Woe=10 (kJ) oxide wt % cations activity sd % SiO2 33.76 2.629 phl 0.0088 0.00332 38 TiO2 3.16 0.185 ann 0.090 0.0165 18 Al2O3 19.71 1.810 east 0.015 0.0051 34 Cr2O3 0.24 0.015 Fe2O3 0.00 0.000

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241 FeO 24.02 1.565 MnO 0.13 0.009 MgO 5.28 0.612 CaO 0.13 0.011 Na2O 0.40 0.061 K2O 8.20 0.816 Totals 95.04 7.712 ________________________________________________________________________ ________ fsp plag Holland & Powell 1992 model 1 Ferric from: all ferric plag is C1 structure oxide wt % cations activity sd % SiO2 59.66 2.679 an 0.47 0.0266 6 TiO2 0.05 0.002 ab 0.65 0.0324 5 Al2O3 24.50 1.297 Cr2O3 0.08 0.003 Fe2O3 0.29 0.010 FeO 0.00 0.000 MnO 0.05 0.002 MgO 0.13 0.009 CaO 7.23 0.348 Na2O 7.21 0.627 K2O 0.09 0.005 Totals 99.31 4.981 ________________________________________________________________________ ________ cd rim 2-site Mg-Fe-Mn mixing, anhydrous basis Ferric from: Cation Sum = 11 for 18 oxygens. Max Ratio = 0.2 WFeMg=1.5, WMgMn=1.5, WFeMn=0 (kJ) oxide wt % cations activity sd % SiO2 49.17 5.063 crd 0.31 0.0313 10 TiO2 0.12 0.009 fcrd 0.23 0.0249 11 Al2O3 31.48 3.821 mncrd Cr2O3 0.15 0.012 Fe2O3 1.33 0.103

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242 FeO 9.94 0.856 MnO 0.17 0.014 MgO 6.67 1.023 CaO 0.14 0.015 Na2O 0.35 0.070 K2O 0.09 0.011 Totals 99.61 11.000 THERMOCALC OUTPUT: Average PT calculation THERMOCALC 3.21 An independent set of react ions has been calculated Activities and their uncertainties py gr alm phl ann east an a 0.00250 4.80e-5 0.470 0.00880 0.0900 0.0150 0.470 sd(a)/a 0.67228 0.83577 0.15000 0.56674 0.29400 0.51082 0.10000 crd fcrd q sill H2O a 0.310 0.230 1.00 1.00 1.00 sd(a)/a 0.14283 0.17787 0 0 Independent set of reactions 1) gr + q + 2sill = 3an 2) 2py + 5q + 4sill = 3crd 3) 2alm + 5q + 4sill = 3fcrd 4) py + east + 3q = phl + crd 5) alm + east + 3q = ann + crd Calculations for the inde pendent set of reactions (for a(H2O) = 1.0) P(T) sd(P) a sd(a) b c ln_K sd(ln_K) 1 4.7 1.42 25.36 0.59 -0.11333 5.322 7.679 0.888 2 4.2 1.16 -60.65 1.05 -0.05391 10.361 8.469 1.411 3 4.9 0.48 74.95 1.08 -0.10307 11.164 -2.899 0.612 4 4.2 2.57 -40.88 0.52 -0.00976 3.411 4.287 1.027 5 2.3 1.46 3.48 1.06 -0.02315 3.708 1.376 0.625 Average PT (for a(H2O) = 1.0) Single end-member diagnostic information

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243 avP, avT, sd's, cor, fit are result of doubling the uncertainty on ln a : a ln a suspect if any are v different from lsq values. e* are ln a residuals normalised to ln a uncertainties : large absolute values, say >2.5, point to suspect info. hat are the diagonal elements of the hat matrix : large values, say >0.45, poi nt to influential data. For 95% confidence, f it (= sd(fit)) < 1.61 however a larger value may be OK look at the diagnostics! avP sd avT sd cor fit lsq 3.8 0.9 657 88 0.901 1.08 P sd(P) T sd(T) cor fit e* hat py 2.71 1.32 560 124 0.958 0.90 0.73 0.75 gr 3.76 0.91 657 88 0.895 1.08 0.08 0.07 alm 3.80 0.88 648 88 0.845 1.05 -0.25 0.23 phl 3.87 0.84 661 82 0.899 0.95 -0.90 0.04 ann 4.00 0.86 673 83 0.903 0.81 0.95 0.03 east 3.88 0.85 667 83 0.903 0.97 -0.83 0.03 an 3.76 0.90 657 88 0.899 1.08 -0.03 0.01 crd 4.05 1.01 681 95 0.924 1.03 0.34 0.14 fcrd 3.80 0.85 686 96 0.810 1.01 -0.42 0.63 q 3.77 0.90 657 88 0.901 1.08 0 0 sill 3.77 0.90 657 88 0.901 1.08 0 0 T = 657C, sd = 88, P = 3.8 kbars, sd = 0.9, cor = 0.901, sigfit = 1.08 THERMOCALC 3.21 Pressures calculated over the 750-900C range An independent set of react ions has been calculated Activities and their uncertainties py gr alm phl ann east an a 0.00250 4.80e-5 0.470 0.00880 0.0900 0.0150 0.470 sd(a)/a 0.67228 0.83577 0.15000 0.56674 0.29400 0.51082 0.10000 crd fcrd q sill H2O a 0.310 0.230 1.00 1.00 1.00 sd(a)/a 0.14283 0.17787 0 0 Independent set of reactions 1) gr + q + 2sill = 3an 2) 2py + 5q + 4sill = 3crd

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244 3) 2alm + 5q + 4sill = 3fcrd 4) py + east + 3q = phl + crd 5) alm + east + 3q = ann + crd Calculations for the inde pendent set of reactions at T = 825C (for a(H2O) = 1.0) P(T) sd(P) a sd(a) b c ln_K sd(ln_K) 1 5.4 1.53 25.35 0.59 -0.11329 5.314 7.679 0.888 2 4.1 1.26 -58.83 1.05 -0.05522 10.261 8.469 1.411 3 5.8 0.51 76.45 1.08 -0.10410 11.067 -2.899 0.612 4 3.7 2.78 -40.04 0.52 -0.01042 3.377 4.287 1.027 5 2.5 1.58 4.16 1.06 -0.02367 3.677 1.376 0.625 Average pressures (for a(H2O) = 1.0) Single end-member diagnostic information av, sd, fit are result of doubli ng the uncertainty on ln a : a ln a suspect if any are v different from lsq values. e* are ln a residuals normalised to ln a uncertainties : large absolute values, say >2.5, point to suspect info. hat are the diagonal elements of the hat matrix : large values, say >0.45, poi nt to influential data. For 95% confidence, f it (= sd(fit)) < 1.54; however a larger value may be OK look at the diagnostics! av sd fit lsq 5.33 0.59 1.31 P sd fit e* hat a(obs) a(calc) py 5.42 0.59 1.27 0.7 0.09 0.00250 0.00403 gr 5.32 0.60 1.31 -0.1 0.07 4.80e-5 4.42e-5 alm 5.32 0.73 1.31 -0.0 0.21 0.470 0.468 phl 5.42 0.55 1.21 1.0 0.04 0.00880 0.0158 ann 5.46 0.48 1.06 -1.2 0.02 0.0900 0.0632 east 5.37 0.53 1.19 1.2 0.01 0.0150 0.0272 an 5.33 0.59 1.31 0.0 0.01 0.470 0.472 crd 5.50 0.52 1.12 -0.9 0.04 0.310 0.271 fcrd 4.74 0.65 1.08 1.2 0.48 0.230 0.282 q 5.33 0.59 1.31 0 0 1.00 1.00 sill 5.33 0.59 1.31 0 0 1.00 1.00 TC 750 770 790 810 830 850 870 890 900 av P 4.6 4.8 5.0 5.1 5.3 5.5 5.7 5.9 6.0 sd 0.44 0.47 0.51 0.55 0.59 0.63 0.67 0.71 0.73 sigfit 1.1 1.1 1.2 1.3 1.3 1.4 1.4 1.5 1.5

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APPENDIX F GEOLOGIC MAP OF THE ANACONDA METAMORPHIC CORE COMPLEX LOWER PLATE

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246 Object F-1. Geologic map of the An aconda metamorphic core complex lowe r plate (AMCCfinalmap.pdf, 22 mb).

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255 Lageson, D. R., Schmitt, J. G., Horton, B. K., Kalakay, T. J., and Burton, B. R., 2001, Influence of Late Cretaceous magmatism on the Sevier orogenic wedge, western Montana: Geology, v. 29, p. 723-726. Lavier, L. L., W. R. Buck & A. B. N. Poliakov, 1999, A self-consistent rolling-hinge model for the evolution of large-offset lo w-angle normal faults: Geology, v. 27, p. 1127-1130. Lewis, R. S., 1990, Geologic map of the Dickie Peak Quadrangle, Deer Lodge and Silver Bow Counties, Montana: Montana Bureau of Mines and Geology, geologic map 51. Lewis, R., 1998, Geologic map of the Butte 1 by 2 degree quadrangle, scale 1:250,000: Montana Bureau of Mines and Geology, Butte, open file report MBMG 363. Lidke, D. J., and Wallace, C. A., 1993, Rocks and structure of the north-central part of the Anaconda Range, Deer Lodge and Granite Counties, Montana: Unites States Geological Survey Bulletin, 31 p., map scale 1:24000. Lister, G. S., and Baldwin, S. L., 1996, Modelin g the effect of arbitr ary P-T-t histories on argon diffusion in minerals using the MacA rgon program for the Apple Macintosh: Tectonophysics v. 253, p. 83-109. Lister, G. S., and Baldwin, S. L., 1993, Plut onism and the origin of metamorphic core complexes: Geology, v. 21, p. 607-610. Lister, G. S., and Baldwin, S. L., 1989, The origin of metamorphic core complexes and detachment faults formed during Tertiary continental extension in the northern Colorado River region, U.S.A.: Journal of St ructural Geology, v. 11, p. 65-94. Lister, G. S., and Davis, G. A., 1989, The origin of metamorphic core complexes and detachment faults formed during Tertiary continental extension in the northern Colorado River region, U.S.A.: Journal of St ructural Geology, v. 11, p. 65-94. Liu, M., 2001, Cenozoic extension and magmatis m in the North American Cordillera: the role of gravitational collapse : Tectonophysics, v. 342, 407-433. Lonn, J. D., McDonald, C., Lewis, R. S., Kalaka y, T. J., ONeill, J. M., Berg, R. B., and Hargrave, P., 2003, Preliminary geologi c map of the Philipsburg 30x 60 quadrangle, western Montana: Montana Bureau of Mines and Geology Open file Report MBMG 483, scale: 1:100,000. Ludwig K. R., 2004, Users manual for ISOPLO T, a geochemical toolkit for Microsoft Excel (v. 3.09a).

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261 BIOGRAPHICAL SKETCH Warren Grice was born in Macon, GA, in 1979. He graduated from Tattnall Square Academy high school (Macon, GA) in 1998. He received a Bachelor of Science degree in geology (with honors) from Montana State University (Bozeman, MT) in 2002. While attending graduate school at the University of Florida (Gai nesville, FL) Warren served as a teaching assistant for several courses in the Department of Geological Sciences and as a research assistant for Dr. David A. Foster. He will receive a masters degree in geology (with high honors) from the University of Florida in August of 2006. Warren will continue to pursue his interests in geology by enrolling in a PhD program in geology or with a career in the industry.


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Title: Exhumation and cooling history of the Middle Eocene Anaconda metamorphic core complex, western Montana
Physical Description: Mixed Material
Copyright Date: 2008

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EXHUMATION AND COOLING HISTORY OF THE MIDDLE EOCENE
ANACONDA METAMORPHIC CORE COMPLEX,
WESTERN MONTANA













By

WARREN CALHOUN GRICE, JR.


A THESIS PRESENTED TO THE GRADUATE SCHOOL
OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT
OF THE REQUIREMENTS FOR THE DEGREE OF
MASTER OF SCIENCE

UNIVERSITY OF FLORIDA


2006

































Copyright 2006

by

WARREN CALHOUN GRICE, JR.

































Dedication: To my family for their never ending support and inspiration.















ACKNOWLEDGMENTS

I would like to thank David Foster, my advisor, for his guidance and patience while

working with me on this study over the past few years. I thank Tom Kalakay for

introducing me to geology and for instilling in me great passion and motivation for the

subject; I also thank Kalakay for his essential guidance in the field during the 2004 field

season. I am greatly appreciative of Jim Vogl, Phil Neuoff, Paul Mueller, Matt Smith,

and Mike Perfit for their assistance and guidance with several aspects of the study. I

thank George Kamenov, Sam Coyner, and Tom Beasley who helped make the analytical

aspects of the study possible. I thank Dan Gorman for help with sample preparation. I

would also like to thank Michaela Speirs, Heather Bleick, Syklar Pauli, and Mike

McTeauge for their assistance in the field; I also thank McTeague for sharing my joy for

geology over the past several years and for many interesting geologic discussions.

Financial support for the field aspect of this study was supplied by grants obtained

by author from the Tobacco Root Geological Society, the Geological Society of America,

and the Belt Association. Funding for the analytical portion of the study was supplied by

a USGS EDMAP grant obtained by Tom Kalakay and by grants obtained by David

Foster. The study would not have been possible without this funding.
















TABLE OF CONTENTS



A C K N O W L E D G M E N T S ................................................................................................. iv

LIST OF TABLES ........... .... ......... ........... ........... .... ......... ............ .. ix

LIST OF FIGURES ............................... ... ...... ... ................. .x

ABSTRACT .............. .................. .......... .............. xiv

CHAPTERS

1 IN T R O D U C T IO N ............................................................................. .............. ...

2 REGIONAL GEOLOGIC BACKGROUND ........................................ ....................9

P re-M esozoic H history ....................................................... 9
M esozoic H history ........ ....... .... ................................................. ..... ........ ...... 13
Crustal Shortening along the Lewis and Clark Line during the Late
Cretaceous .................. .. .................. .......... 14
Late Cretaceous M agmatism ........... .......... ..... .. ..... ............... 16
High Grade Metamorphism during the Cretaceous .....................................19
E ocene H history ................................ ...... .... ...... ...... ..... .................20
Regional Transtension and Large-scale Crustal Extension along the Lewis
and Clark Line................ ......... ......... .........21

3 THE ANACONDA METAMORPHIC CORE COMPLEX ....................................23

Structural-metamorphic Domains of the Anaconda Metamorphic Core Complex: ...23
T he L ow er P late ........................................ ... .... ........ .... .. ... 23
The D etachm ent Fault Zone ........................................... ......................... 29
T he U pper Plate ................................................................................................... 33
Description of the Lake of the Isle Shear Zone..................................... ................. 35
Description and Distribution of Metamorphic and Structural fabrics.................36
Metamorphic foliations and gneissic banding.................. ................36
M esoscopic-scale folds ........................................ .......................... 41
M esoscopic-scale boudins.................................... ..................................... 42
Shear sen se ................ ............ ........... ...... .... ......... ................. 4 5
Description and Distribution of Metamorphic Phase Assemblages and
T ex tu re s ...................................... ............................................. 4 5









Metamorphic phase assemblages and textures in the meta-Greyson
Form action ............... .................................................. ..... ... ......... 46
Relict and fresh kyanite in the upper Meta-Ravalli Group ........................50
Spatial Relationship Between the Lake of the Isle Shear Zone and Late
C retaceous Intrusions .............................................................. ....... ... .... ..... .. 53

4 U-PB ZIRCON GEOCHRONOLOGY .......................................... ...............58

P purpose and Strategy .............. .. ..................................................... ................ .. 58
Relevant Previous U-Pb Zircon Geochronology ................................................59
U-Pb Zircon Geochronology Results.............................. ........................ 59
WG04-114 (Storm Lake Stock Granodiorite)...............................................60
Ug-1 (The Deformed Quartz Diorite Sill) ..................................61
WG05-02 (Leucosome from the Meta-Greyson Paragneiss)...........................70

5 THERM OBAROM ETRY ................................................ .............................. 76

Purpose and Strategy .............. ...... ........ ...... ... ... ..... ............ 76
ME-231 (Migmatitic Meta-Greyson Formation Paragneiss)............................... 77
Relevant Previous Pressure-temperature Constraints.................. ....... .........78
Therm obarom etry R results ................................................ .............................. 80
Electron M icroprobe Analyses ...................................................................... 80
G arnet ............................... .................................. .........80
B io tite ..................................................................................................... 8 6
Plagioclase............................................. 86
Cordierite ............... ....................................... 86
Mineral End Member Activity Calculations made using AX .............................87
Pressure-temperature Estimates using THERMOCALC ....................................87
Pressure-temperature Estimates using Geothermobarometry (GTB)..................93

6 40AR/39AR THERMOCHRONOLOGY ................ .............................................95

Purpose and Strategy ........................................ ................. .... ....... 95
P previous T herm ochronology ........................................................... .....................96
40A r/39A r Therm ochronology Results ........................................... .................... 104
Argon Closure Temperature ............. ............................................ ....... 105
B io tite ...................... .. ............. .. ..................................................1 2 2
M uscovite ......................................... .................. .. .... ........ 125
Hornblende ......................................... .................. .... ........ 128
K -feldspar ................................... ........................... 13 1
40Ar/39Ar Thermochronology From Outside the Lower Plate Transect .................133

7 D ISC U S SIO N ................................................................................ ............... 136

Origin of the Lake of the Isle Shear Zone ..............................................136
A g e C o n strain ts ........................................................................................... 13 7
Pressure-tem perature H history .......................................................................... .140
Kinematic Interpretation ....................................................... 144









Strain localization and ductile attenuation of metasedimentary strata in
th e L IS Z ............................................................................. 14 5
Strain classification .......................... ...... ................ ...... .. .. ...... .. 146
Structural Interpretation ....................... ................. .................. ... 146
Eocene Exhumation and Cooling History of the Anaconda Metamorphic Core
Complex defined by 40Ar/39Ar Thermochronology .............................................147
Low er Plate Cooling H istory.................................... ............................ ....... 148
Constraints on the Timing of the Onset of Extension ............ ... ................154
Constraints on the Duration of Extension ............................... ................159
Constraints on the Detachm ent Slip Rate.................................. ... ................ 161
Constraints on the Original Detachment Geometry .......................................169
Magnitude of Offset on the Detachment ................................ ...............177

8 C O N C L U SIO N S ........................................................................... ............ ... 183

Age and Pressure-temperature Constraints on High Grade Metamorphism in the
ACC lower plate: Origin of the Lake of the Isle Shear Zone..............................183
Exhumation and Cooling History of the Middle Eocene Anaconda Metamorphic
C ore C om plex : .................................................................................... 183
R regional Tectonic Context ......................................................... .............. 188

APPENDIX

A M E T H O D O L O G IE S ..................................................................... .................... 19 1

Field Mapping and Sampling Methods................. ...............................................191
40Ar/39Ar Thermochronology M ethods................................. ....................... 192
Therm obarom etry ........................ ........ .. ........ .. ............. 195
U -P b G eochronology ......... ......... .. ............. .. .. ..................... ........................ 196
M ajor and Trace Elem ent Geochem istry.............................................................. 203

B ACC LOWER PLATE LITHOLOGIC UNIT DESCRIPTIONS AND
D ISTR IB U TIO N ................................................. .. ........ .... .............. .. 205

G general Statem ent.......... ... .................................................... .. ...... .. ...... .. 205
Belt Supergroup-equivalent Metasedimentary Strata.............................206
Middle Cambrian-equivalent Metasedimentary Strata............................................210
L ate C retaceous Intrusions ............................................... ............................ 212
The Storm L ake Stock .............................. .............. .... ....................... 212
The Deformed Quartz Diorite Sill ..........................................................216
E ocene Intru sion s........ ....................................................................... ....... .. ..... .. 2 17

C DESCRIPTION OF SAMPLE ME-231 ....................................... ............... 224

D MAJOR AND TRACE ELEMENT GEOCHEMISTRY .............. ... ...............229

P u rp o se ............................................ ......................... ................ 2 2 9
M ajor and Trace Elem ent Analytical Results..........................................................230









Major Element Analyses ......... ........ ......... ................. 231
Trace Element Analyses ........... .................................. 231

E THERMOCALC AND AX OUTPUT FILES ........................................................ 240

F GEOLOGIC MAP OF THE ANACONDA METAMORPHIC CORE COMPLEX
LOW ER PLATE..................... .......... ...... ...... ...... .............. 245

R E FE R E N C E S C IT E D .......................................................................... ....................247

B IO G R A PH IC A L SK E T C H ........................................ ............................................261












































viii
















LIST OF TABLES


Table page

3-1. Description and distribution of key mineral phases and textures in the
metamorphosed Greyson Formation. ............................................ ............... 54

4-1. U-Pb LA-MC-ICP-MS analytical results for WG04-114 .....................................62

4-2. U-Pb LA-MC-ICP-MS analytical results for Ug-1 .............................................66

4-3. U-Pb LA-MC-ICP-MS analytical results for WG05-02 .......................................68

5-1. Results from electron microprobe mineral analyses. .............................................81

5-2. Averaged ME-231 electron microprobe analyses used in thermobarometry ...........88

6-1. Summary of relevant previous thermochronology................................. ..........98

6-2. Summary of 40Ar/39Ar thermochronology from the Anaconda metamorphic core
com plex .......................................................................... 106

D-1. Results from major and trace element analyses ............................................. 232
















LIST OF FIGURES


Figure page

1-1. Regional geologic map of the North American Cordillera. Emphasis is placed
on major igneous and tectonic features. .................... ............ ... ............ 2

1-2. Tectonic map of northwestern United States and southern Canada showing
major structures related to Tertiary extension.................................................... 4

2-1. Regional geologic map of the major Archean basement provinces and
Proterozoic suture zones of the North American Cordillera in the western United
States and southern Canada ........................................................ ............. 10

2-2. Simplified tectonic map of Belt Basin in the Mesoproterozoic. Map shows the
major normal faults active during deposition of the Belt-Purcell Supergroup. ......12

2-3. Tectonic map of the North American Cordillera of western Montana, northern
Idaho, northeastern Washington, and southern Canada during the Late
C retaceous............................................................... .... ..... ........ 15

2-4. Geologic map of western Montana, south of the Lewis and Clark line ................ 18

3-1. Geological Map of the Anaconda metamorphic core complex (ACC) in western
M ontana....................................................... ................... .... ....... ....... 25

3-2. Geologic cross section of the Anaconda metamorphic core complex, western
M ontana....................................................... ................... .... ....... ....... 26

3-3. Photomosaic of the Anaconda metamorphic core complex.................................32

3-4. Greenschist facies m ylonites .................................. ............... ............... 33

3-5. Outcrop photos of late listric-shaped brittle normal faults in the greenschist
faces m ylonite zone ....................... .................. .................... .. ...... 34

3-6. Outcrop photograph showing the strong gneissic banding of the metamorphosed
Greyson Form ation in the LISZ. ........................................ .......................... 39

3-7. Outcrop photographs of gneissic banding commonly observed in the LISZ...........40

3-8. Photomicrograph of the Missoula Group-equivalent calc-silicate paragneiss.........41









3-9. Outcrop photograph of the deformed quartz diorite sill in the center of the
L IS Z .. .................................................................................... 4 3

3-10. Outcrop photographs of mesoscopic-scale folds found in Belt equivalent
metasedimentary strata deformed in the LISZ.. ............................ ..... .......... 44

3-11. Outcrop photographs of mesoscopic-scale folds found in Belt-equivalent
metasedimentary strata deformed in the LISZ.. ............................ ..... .......... 46

3-12. Outcrop photographs of mesoscopic boudins in the LISZ...................................47

3-13. Outcrop photographs of shear sense indicators in the meta-Greyson Formation
deform ed in the L ISZ .. ......................... .................... ............... .... ...... ...... 49

3-14. Outcrop photographs of meta-Greyson migmatitic paragneiss showing granitic
leucosome commonly found in pressure shadows between quartzite boudins........51

3-15. Photomicrographs of thin sections of the meta-Greyson migmatitic paragneiss
exhibiting granitic leucosome (sample ME-231)........... .................................... 52

3-16. Simplified geologic sketch map of the ACC lower plate exposed in the current
study area showing the metamorphic mineral/textural zones 1, 2, and 3 mapped
in the meta-Greyson Formation adjacent to the deformed quartz diorite sill...........55

3-17. Kyanite pseudomorphs in the upper meta-Ravalli Group. A) Outcrop
photograph of kyanite pseudomorphs................ .. ......... ... ............... 56

3-18. Photomicrograph showing fresh kyanite in the upper meta-Ravalli Group.............57

4-1. Tera-Wasserburg plot and 206Pb / 238U weighted mean zircon age plot for Storm
Lake Stock granodiorite sample W G04-114..........................................................70

4-2. Tera-Wasserburg plot and 206Pb / 238U weighted mean zircon age plot for the
deformed quartz diorite sill sample Ug-1 ....... .............................................. 71

4-3. Conventional U-Pb concordia plots for zircons for leucosome sample WG05-02
collected from the meta-Greyson migmatitic paragneiss in the central LISZ..........74

4-4. 207Pb/206Pb age density plot for the WG05-02 leucosome zircons. 207Pb/206Pb
ages range from -1525-3004 M a.. ............................ ............................................75

5-1. Phase diagrams showing PT estimates for ME-231, migmatitic Greyson
paragneiss from the Lake of the Isle shear zone................................................91

5-2. Phase diagram showing phase equilibria used by THERMOCALC to calculate
the average PT estimate for sample M E-231 ........................................................93

6-1. Geologic map of Anaconda metamorphic core complex (ACC), western
Montana and vicinity showing location of previous thermochronology..............102









6-2. Geologic sketch map showing mineral cooling ages from samples collected
along the ACC lower plate transect using 40Ar39Ar thermochronology...............109

6-3. 40Ar/39Ar mineral age spectra obtained for samples collected from the Anaconda
metamorphic core complex using 40Ar/39Ar thermochronology. ...........................110

7-1. Phase diagram showing the Late Cretaceous PT history of the Lake of the Isle
shear zone (LISZ) .......................................... ............... .. ........ .... 142

7-2. Geologic sketch map showing the mineral cooling ages obtained by 40Ar39Ar
thermochronology from samples collected along the ACC lower plate transect in
th is stu dy .. ...............................................................................14 9

7-3. Mica cooling age contour map constructed from biotite and muscovite cooling
ages obtained 40Ar39Ar thermochronology from samples collected along the
A CC low er plate transect in this study................................................................ 150

7-4. Temperature-time cooling diagram showing the cooling history of the ACC
low er p late ................................................................................15 1

7-5. Age vs. distance in slip direction diagram and lower plate transect sketch map.. .159

7-6. Slip rate calculations.. .............................. ... ........................................166

7-7. Paleoisotherm contour map. Paleoisotherms refer to estimated temperatures of
the ACC lower plate directly beneath the detachment at the onset of extension at
-53 M a. .......... ......................................... ..... .......... ....... ......... 174

7-8. Distance in slip direction vs. paleodepth diagram showing geometries for the
detachment at the onset of extension at -53 M a. ................................................ 175

7-9. Distance in slip direction vs. paleodepth diagram showing magnitude of offset
on the detachment with variable slip rate from -53-40 Ma with a geothermal
gradient of 35 C/km .......................................... ................ ............ 181

7-10. Geologic map of the ACC showing structural pinning points used to constrain
the maximum amount of displacement facilitated by the Anaconda detachment.. 182

D-1. Major element Harker variation diagrams for samples Ug-1, WG04-114, and
D F 02 -114 ........................................................................ 2 36

D-2. REE trend and trace element spider diagrams for samples Ug-1, WG04-114, and
DF02-114.. .................................... .................................. .......... 239
















OBJECT LIST


Object


F-1. Geologic map of the Anaconda metamorphic core complex lower plate
(AM CCfinalmap.pdf, 22 mb) ........... ....................... ................. 246


page















Abstract of Thesis Presented to the Graduate School
of the University of Florida in Partial Fulfillment of the
Requirements for the Degree of Master of Science

EXHUMATION AND COOLING HISTORY OF THE MIDDLE EOCENE
ANACONDA METAMORPHIC CORE COMPLEX, WESTERN MONTANA

By

Warren Calhoun Grice, Jr.

August 2006

Chair: David A. Foster
Major: Department: Geological Sciences

New 40Ar/39Ar thermochronology, U-Pb geochronology, and thermobarometry

define the tectonic exhumation and cooling history of the Middle Eocene Anaconda

metamorphic core complex (ACC) of western Montana. Mica 40Ar/39Ar cooling ages

obtained from the ACC lower plate footballl): (1) constrain the age of the onset of

extension in the ACC to -53 Ma; (2) constrain the duration of extension to at least -53-

39 Ma; (3) define a lateral cooling age gradient, where mica cooling ages decrease to the

ESE across the ACC lower plate confirming top-to-the-ESE directed unroofing of the

lower plate; (4) constrain the original geometry of the Anaconda detachment at -53 Ma

to a listric-shaped normal fault comprised of a steep portion (-54-70) in the upper brittle

crust and a sub-horizontal portion (-7-12) in the middle crust; and (5) constrain the rate

of slip on the detachment to -0.9 km/Myr during the time interval of -53-39 Ma.

Reconstruction of similar Late Cretaceous granodiorite plutons in the detached ACC

upper plate (hanging wall) and lower plate indicate -25-28 km of (horizontal)









displacement occurred on the detachment during the Eocene. Thermobarometry and U-

Pb zircon geochronology from the Lake of the Isle shear zone (LISZ, named herein),

structurally beneath the detachment, show that uppermost-amphibolite facies

metamorphism (at -3.2-5.3 kbar and -750-8500C) ended in the Late Cretaceous (at -75-

74 Ma). Pressure constrains from the thermobarometry indicate the lower plate was

exhumed from a maximum crustal depth of-10-16 km.

The results from this study show the exhumation and cooling histories of the ACC

and the Bitterroot metamorphic core complex (BCC), located -70 km west of the ACC,

are remarkably similar. Therefore, the ACC and BCC represent one continuous and

integrated extensional system that accommodated large-scale extension in easternmost

Idaho and western Montana during the Eocene. Extension in the ACC and BCC was

linked to regional dextral transtension along the Lewis and Clark Line (LCL), a major

strike-slip fault zone located to the north of these metamorphic core complexes.

Extension in the ACC and BCC, combined with extension in several other metamorphic

core complexes north of the LCL, is responsible for the initial collapse of the previously

thickened Sevier hinterland beginning in the early to middle Eocene immediately

following (-1-3 Ma) the end of crustal shortening in the foreland fold-and-thrust-belt to

the east. Regional dextral transtension, large-scale extension, and exhumation of

metamorphic core complexes in the northwestern United States and southern British

Columbia were likely driven by traction caused by increased obliquity of convergence at

the western margin of the North American Plate beginning in the early to middle Eocene.














CHAPTER 1
INTRODUCTION

Metamorphic core complexes are regional extensional structures characterized by

large-scale crustal extension and tectonic exhumation of mid-crustal metamorphic and

plutonic rocks along low-angle normal brittle-ductile detachment fault zones (e.g., Davis

and Coney, 1979; Coney, 1980). More than thirty metamorphic core complexes have

now been documented throughout the North American Cordillera and comprise a

relatively narrow and sinuous belt stretching from southern British Columbia into

northern Mexico (Coney, 1980; Armstrong and Ward, 1991). These metamorphic core

complexes formed during early to middle Tertiary time following a long period of crustal

accretion, shortening, and thickening caused by convergence along the western margin of

North American during the Mesozoic (Burchfiel et al., 1992; Wernicke, 1992). With the

exception of the southernmost metamorphic core complexes, all the core complexes

formed in the previously thickened Mesozoic Sevier hinterland, west of the Sevier thin-

skinned foreland fold-and-thrust belt and Laramide-style basement-involved foreland

uplifts (Fig. 1-1; Coney and Harms, 1984; Wernicke, 1992). Large-scale extension in

metamorphic core complexes occurred before widespread Basin-and-Range-style

extension affected the North American Cordillera (Liu, 2001). Therefore, metamorphic

core complexes are responsible for the initial extensional collapse of the previously

thickened Cordilleran crust during the Tertiary (Coney and Harms, 1984; Wernicke,

1992; Foster et al., 2001; Vanderhaeghe et al., 2003).


















































Figure 1-1. Regional geologic map of the North American Cordillera. Emphasis is
placed on major igneous and tectonic features. Note the location of the
Cordilleran metamorphic core complexes located west of the Sevier fold-and-
thrust-belt. The Anaconda metamorphic core complex (ACC) is located south
of the Lewis and Clark line in western Montana. Note the Lewis and Clark
Fault Zone = the Lewis and Clark Line. (Modified from Coney, 1980, and
Foster et al., 2006a).









Large-scale extension and exhumation of mid-crustal metamorphic and plutonic

rocks in metamorphic core complexes in the northwestern United States was linked to

regional dextral transtension along the Lewis and Clark line (LCL) during the Eocene

(Doughty and Sheriff, 1992; Yin and Oertel, 1995; Foster et al., 2006a). The LCL is a

northwest-west-trending -40-80 km wide zone of steeply-dipping strike-slip, oblique-

slip, and dip-slip faults that reaches from northeastern Washington to west-central

Montana. Many faults of the LCL accommodated repeated displacement during multiple

tectonic reactivations since the inception of the line in the Mesoproterozoic or earlier

(Wallace et al., 1990; Sears and Hendrix, 2004). During the Late Cretaceous to early

Eocene (-88-55 Ma) the LCL accommodated sinistral displacement and regional

transpression related to crustal shortening and thickening in the Sevier fold-and-thrust

belt of southern Alberta and western Montana. Beginning in the early Eocene (-54-52

Ma), the LCL was reactivated as a regional dextral transfer structure related to large-scale

extension and exhumation of several metamorphic core complexes in the Sevier

hinterland, west of the fold-and-thrust belt (Fig. 1-1 and 1-2, Sears and Hendrix, 2004;

Foster et al., 2006a).

South of the LCL, in eastern Idaho and western Montana, large-scale Eocene

extension may have been accommodated by exhumation of both the Bitterroot and

Anaconda metamorphic core complexes (Fig. 1-2, Foster et al., 2006a). The timing and

kinematics of Eocene extension in the Bitterroot metamorphic core complex (BCC), the

more westerly of the two core complexes, is n well constrained (e.g., Hyndman, 1980;

House and Applegate, 1993; House and Hodges, 1994; Foster and Fanning, 1997; Foster,

2000; Foster et al., 2001; House et al., 2002; Foster and Raza, 2002). Extension in the









BCC began at -53 Ma and continued to -25 Ma (based on a large U-Pb geochronology,

40Ar/39Ar, and fission track thermochronological data set from the exhumed football,


Figure 1-2. Tectonic map of northwestern United States and southern Canada showing
major structures related to Tertiary extension. The Bitterroot and Anaconda
metamorphic core complexes are located south of the Lewis and Clark Line
(LCL). The LCL was active as a regional dextral transfer structure between
extension north and south of the line during the Eocene (Modified from Foster
et al., 2006a).









Foster et al., 2001; Foster and Raza, 2002; Foster et al., 2006a). During this time, the

BCC lower plate was exhumed from beneath a single east-dipping upper-amphibolite

facies mylonitic shear zone and overprinting brittle normal detachment fault system by

top-to-the-east-southeast directed (110-100) detachment of the hanging wall

(constrained by a large thermochronology data set from the football and kinematic

indicators and mineral stretching lineations from the mylonites, Foster, 2000; Foster, et

al., 2001; House et al., 2002; Foster and Raza, 2002). Early extension in the BCC was

accompanied by upper-amphibolite facies metamorphism at -6-8 kbar and -650-750C

and localized decompressional anatexis (partial melting) based on thermobarometry and

U-Pb geochronology from migmatites directly beneath the mylonitic shear zone in the

eastern parts of the football (House et al., 1997; Foster et al., 2001; Foster and Raza,

2002). The thermobarometric data show that the eastern BCC was exhumed from lower

to mid-crustal depths of -20-25 km (Foster and Raza, 2002; Foster et al., 2006a). Total

(horizontal) displacement on the detachment in the BCC is estimated to be -40-50 km,

based on reconstruction of Cretaceous dioritic plutons in the detached hanging wall with

similar dioritic plutons in western football (Foster et al., 2006a).

Extension in the Anaconda metamorphic core complex (ACC, O'Neill, et al., 2002;

Kalakay and Lonn, 2002; Kalakay et al., 2003; O'Neill et al., 2004), located -70-80 km

east of the BCC, is not well constrained, but apparently remarkably similar (Fig. 1-2).

Limited U-Pb geochronological and thermochronological data suggest extension in the

ACC began at -53 Ma, coincident with the onset of extension in the BCC, and continued

until at least -47 Ma (O'Neill et al., 2004; Foster et al., 2006a). As in the BCC, the ACC

lower plate footballl) appears to have been exhumed along a single east-dipping









mylonitic shear zone and overprinting brittle detachment fault system which now bounds

its eastern side (Kalakay et al., 2003; O'Neill et al., 2004). Kinematic indicators and

mineral stretching lineations from the mylonites suggest top-to-the-east-southeast (102-

1080) directed detachment of the ACC upper plate (hanging wall) exhumed the ACC

lower plate (Kalakay et al., 2003). However, a rather large thermochronological data set

is needed from across the ACC to definitely prove east-southeast directed detachment of

the upper along an east-dipping detachment system (e.g., Foster and John, 1999).

Metamorphosed Lower Belt Supergroup pelitic strata exposed directly beneath the

mylonitic shear zone and brittle detachment system in the ACC are migmatitic and show

evidence for upper-amphibolite facies metamorphism and anatexis during high

temperature ductile deformation (T. Kalakay, per comm.). However, the metamorphic

grade, age, and relationship of these migmatites to Eocene extension in the ACC have not

been established. Geochronological and thermobarometric data are needed from the

lower plate migmatites. Within these data, the amount of exhumation in the ACC can not

be quantified. In addition, the total amount of displacement facilitated by the ACC

detachment during the Eocene has not been established.

The apparent similarities in the timing and kinematics of Eocene extension in the

BCC and ACC has led some workers to propose the two metamorphic core complexes

were part of a continuous integrated extensional system that together accommodated

large-scale extension south of the LCL during the Eocene (Doughty and Sheriff, 1992;

O'Neill et al., 2004, Foster et al., 2006a). Following this idea, the BCC and ACC would

have been exhumed along separate synthetic (parallel) east-dipping mylonitic shear zones

and brittle detachment systems that would have been active at the exact same time. In









this tectonic model a "nested" more shallow position for ACC mylonitic shear zone and

brittle detachment is proposed (D. Foster, per. comm.; O'Neill et al., 2004, their Fig. 8).

The primary aim of this study is to test the idea that the BCC and ACC

metamorphic core complexes represent a single integrated extensional system responsible

for large-scale Eocene extension south of the LCL as has been suggested. In order to test

the proposed tectonic model, Eocene extension in the ACC must be constrained. In this

study a combination of 40Ar/39Ar thermochronology, U-Pb geochronology,

thermobarometry, major-trace element geochemistry, and field mapping are used to

provide a detailed exhumation and cooling history for the ACC. In particular, this study

will: (1) confirm the age of the onset extension in the ACC; (2) provide constraints on

the duration of extension in the ACC (i.e., did extension in the ACC continue after -47

Ma?); (3) constrain the Late Cretaceous to Late Eocene cooling history for the ACC

lower plate; (4) confirm top-to-the-east-southeast directed tectonic unroofing

(detachment) of the ACC upper plate as in the BCC; (5) constrain on the original

geometry of ACC detachment to confirm an original east-dipping detachment synthetic to

the BCC detachment; (6) constrain the slip rate and magnitude of slip on the ACC

detachment system; (7) constrain metamorphic pressure and temperature conditions and

age of upper-amphibolite facies metamorphism in the ACC lower plate; these

metamorphic pressure constraints are used to (8) constrain the maximum amount of

exhumation in the ACC. The metamorphic pressure constraints are also used to (9)

determine the maximum depth of the ACC detachment in the Eocene to test its proposed

"nested" geometry relative to the BCC detachment.









If the proposed tectonic model for large-scale Eocene extension in the BCC and

ACC is substantiated then the structural style of extension south of the LCL differed

significantly from Eocene extension in and north of the LCL in northern Idaho,

northeastern Washington, and southern British Columbia; as all of these core complexes

were exhumed by paired east and west-dipping detachment systems (e.g., the Shuswap

metamorphic core complex), not by single asymmetric detachments (e.g., the BCC). An

improved and more complete understanding of Eocene extension in the northwestern

Unites States, south of the LCL, is a critical contribution for defining the role of plate

margin forces, mantle upwelling, and extensional (orogenic) collapse which apparently

caused the destruction of an Andean-style orogen that existed in the North American

Cordillera from Cretaceous to early Eocene time (e.g., Coney and Harms, 1984, Foster et

al., 2001; Foster et al., 2006a).














CHAPTER 2
REGIONAL GEOLOGIC BACKGROUND

Pre-Mesozoic History

Archean (>2.5 Ga) and Paleoproterozoic basement underlie western Montana

(Foster et al., 2006b). In northwestern Montana the basement is comprised of the

Medicine Hat Block of the southern Archean Hearne Province, which extends beneath

southernmost British Columbia, Alberta and Saskatchewan. To the south, the underlying

basement consists of the Archean Wyoming Province. These two Archean Provinces are

separated by the linear northeast-trending Great Falls tectonic zone (GFTZ) which

stretches from eastern Idaho to the northeast into Saskatchewan (O'Neill and Lopez,

1985). The GFTZ is mostly Paleoproterozoic in age and has been interpreted to represent

the final suturing of the Archean Wyoming Province and Medicine Hat Block at ca. -1.9-

1.8 Ga (based on U-Pb zircon crystallization ages, Mueller et al., 2003). The Wyoming-

Medicine Hat suture (the GFTZ), along with several other north to northeast-trending

Paleoproterozoic collisional belts within the Canadian shield (e.g., the Trans-Hudson

orogen) were responsible for the assemblage of the Laurentia cratonic core from -2.0-1.8

Ga (Hoffman, 1988).

The Mesoproterozoic Belt Basin is superimposed of the western edge of the

Archean-Paleoproterozoic basement in western Montana and also extends into northern

Idaho, northeastern Washington, and southern Alberta (Fig. 2-1). The Belt Basin

contains a thick (up to -18 km) sequence of fine-grained quartzites, red to green-colored

argillites, argilleous carbonates, and dark-grey well laminated argillites and quartzites




































Figure 2-1. Regional geologic map of the major Archean basement provinces and
Proterozoic suture zones of the North American Cordillera in the western
United States and southern Canada. The map also shows the position of the
Belt-Purcell Basin superimposed (dark dashed outline) on the western edge of
the Precambrian basement provinces in northwestern Washington, northern
Idaho, western Montana, and southern Canada. Darkened shapes represent
subaerial exposes of Precambrian basement. The Sri = 0.706 line represents
the western edge of Precambrian basement (from Foster et al., 2006).

(Winston and Link, 1993). In the northwestern United States this immense sequence of

sedimentary strata is referred to as the Belt Supergroup and in Canada as the Purcell

Supergroup. The timing of deposition of the Belt Supergroup sedimentary strata is now

well constrained to the interval of 1.47-1.40 Ga by U-Pb zircon crystallization ages

from intercalated syn-depositional mafic sills and volcanic tuffs, a large detrital zircon U-

Pb age dataset, and detrital muscovite 40Ar/39Ar cooling ages (Ross and Villeneuve, 2003;

and references therein). The origin and Mesoproterozoic tectonic setting of the Belt









Basin is still controversial (see Winston and Link, 1993 for review). There is increasing

evidence and a general consensus among many workers for a extensional rift-related

intracratonic basin setting for the Belt Basin, where the subsiding basin was in part, if not

fully isolated from an adjacent seaway to the western or northwest (e.g., Winston, 1986;

Winston and Link, 1993; Lyons et al., 2000; Ross and Villeneuve, 2003).

In western-central Montana the Belt Basin narrows significantly to form a

prominent east-west-trending structural embayment referred to as the Helena embayment

(Fig. 2-1 and 2-2). The Helena embayment formed when several major syn-depositional

steep normal faults down-dropped the embayment separate from the rest of the Belt basin

to the west during the Mesoproterozoic (Winston and Link, 1993). The southern margin

of the Helena embayment was bound and down-dropped by the east-west-trending Perry

Line/Willow Creek fault zone. As a result, the Dillon block of the Archean Wyoming

Province was first uplifted along the up-thrown (southern) side of the fault zone (Winston

and Link, 1993). The northern margin of the Helena embayment was bound and down-

dropped along a series of high-angle northwest-west-trending faults which comprised the

eastern continuation of the proto-Lewis and Clark line (see below, e.g., Wallace et al.,

1990; Sears and Hendrix, 2004). Other major faults also active along the eastern segment

of the Lewis and Clark line during the Mesoproterozoic may have included the inferred

east-west-trending Jocko and Garnet lines and the Volcanic Valley fault zone (Fig. 2-2,

Winston and Link, 1993). The geometry of the Helena embayment, as defined by the

Early bounding fault system, had a major influence on subsequent tectonism in western

Montana (especially during the Late Mesozoic to Early Cenozoic, Winston et al., 1986;

Foster et al., 2001, 2006; Sears and Hendrix, 2004).










l lt k IK!' / B W


MEDICINE HAT BLOCK
OF HEARNE PROVINCE
SNorth


WYOMING PROVINCE



ID HELENA EMBAYMENT


S ""- n ODILLON BLOCK OF
'I WYOMING PROVINCE


S Sr S/ Sr 0.706 line --'" ____
1 0 '' ( ,.
0 so 100
118" 1 | ld IHg 11 I1D' I118

Figure 2-2. Simplified tectonic map of Belt Basin in the Mesoproterozoic. Map shows
the major normal faults active during deposition of the Belt-Purcell
Supergroup. The early Lewis and Clark likely included the Osburn Fault
(OF), the Hope Fault (HF), the Jocko Line (JL), and Garnet Line. The JL and
GL are inferred (see Winston and Link, 1993). The Volcano Valley Fault
(VVF) and the Willow Creek Fault Zone (WCF) were responsible for down-
dropping the Helena Embayment during deposition of the Belt-Purcell
Supergroup. The 87Sr/ 86Sr = 0.706 line from Sears and Hendrix (2004) and
represents the western edge of known Precambrian basement. The Great Falls
Tectonic Zone (GFT) was not active during deposition of the Belt-Purcell
Supergroup Mesoproterozoic (Modified from Winston and Link, 1993).

Following the deposition of the Belt Supergroup sedimentary strata, during the

Late Neoproterozoic (ca. -800-700 Ma), a major rift zone developed along the entire

western margin of Laurentia as the proposed pre-Gondwanan supercontinent Rodinia

began to break apart (e.g., Meert and Torsvik, 2003). As a result, the western side of the

Belt Basin was truncated as the adjacent landmass rifted away (the identity of the western

landmass is highly controversial; see Karlstrom et al., 2001 for review). As rifting

continued, a paleo-Pacific Ocean opened and a passive margin sequence was deposited

beginning with the deposition of the Late Neoproterozoic to Early Cambrian-aged









Windermere deep-water sequence along the entire western margin of Laurentia but

mainly outboard (west) of the Belt Basin in western Montana (Burchfiel et al., 1992;

Winston and Link, 1993).

During the Middle Cambrian, a thick miogeoclinal sequence (continental shelf

deposits) was deposited over the eastern Windermere sequence and further inboard onto

the craton during a major transgression (Burchfiel et al., 1992). In western Montana, the

earliest phase of this transgression is represented by the Middle Cambrian Flathead

Sandstone which lies unconformably over upper Belt Supergroup strata in much of

western Montana (Emmons and Calkins, 1913; Winston and Link, 1993).

Mesozoic History

Beginning in the Early Triassic, an active convergent margin developed along the

western margin of Laurentia (Burchfiel et al., 1992). With continued convergence, from

the Early Triassic to Early Cretaceous, several exotic terranes (including a number of

volcanic arc complexes) were accreted to the western edge of Laurentia (e.g., the

Wrangellia terrane and the Wallowa-Seven Devils terrane adjacent to the Western Idaho

Suture, Hamilton, 1978). Despite the major contractional deformation associated with

convergence at this time, the Belt Basin and its Paleozoic miogeoclinal sedimentary

cover sequence remained largely stable until the Late Early to Late Cretaceous (Burchfiel

et al., 1992).

By the Late Early to Late Cretaceous, crustal shortening and thickening related to

the ongoing convergence along the western margin of Laurentia migrated eastward into

present-day northeastern Idaho and western Montana. Thrust faulting began first in

eastern Idaho at -105 Ma and moved progressively eastward into western Montana

before ending at -55 Ma (Hyndman, et al., 1988; Constenius, 1996; Sears and Hendrix,









2004). In western Montana, the east-directed crustal shortening and thickening was

facilitated predominately by low-angle "thin-skinned" basement-detached Sevier-style

thrust faulting superimposed on the Mesoproterozoic Belt Basin (Allmendinger, 1992;

Winston and Link, 1993); this deformation was part of the Sevier orogen which stretches

from British Columbia into southern Nevada (Burchfiel et al., 1992). As a result, many

thin-skinned thrusts in western Montana involve thick sequences of Belt Supergroup and

Paleozoic sedimentary strata thrust eastward over younger Belt, Paleozoic, and Mesozoic

sedimentary strata (Winston, 1986; Foster et al., 2001). In addition, because thrusting

was superimposed on the Belt Basin, a prominent east-directed structural bulge formed in

the thrust belt in the original Helena embayment referred to as the Helena thrust salient

(Fig. 1-2 and 2-3; Winston, 1986).

Crustal Shortening along the Lewis and Clark Line during the Late Cretaceous

Thin-skinned Sevier-style thrusting in western Montana was accommodated by the

eastward displacement of three major thrust slabs (or plates) along the Lewis and Clark

Line (LCL): 1) the Lewis-Eldorado-Hoadley (LEH) slab in northwestern Montana and

the 2) Sapphire and 3) Lombard slabs in southwestern Montana (Constenius, 1996; Sears

and Hendrix, 2004). The LCL, a complex zone of west-to-northwest-trending steeply

dipping strike-slip, dip-slip and oblique slip faults, separates the LEH slab in the north

from the Sapphire and Lombard slabs to the south (Fig. 2-3, Wallace et al., 1990). As

noted above, some faults within the LCL likely originated during subsidence of the Belt

Basin in the Mesoproterozoic. However, several workers show evidence for reactivation

of these old faults as well as inception of younger faults along the LCL during the

Mesozoic (e.g., Harrison et al., 1972; Hyndman et al., 1988; Wallace et al., 1990; Stewart

and Crowell, 1992; Sears and Hendrix, 2004). Most recently, Sears and










































I 118W 114W I
Figure 2-3. Tectonic map of the North American Cordillera of western Montana,
northern Idaho, northeastern Washington, and southern Canada during the
Late Cretaceous. Major thrust slabs in western Montana were comprised of
the Sapphire and Lombard slabs south of the Lewis and Clark Line and the
Lewis-Eldorado-Hoadley slab to the north. Note the eastward budge in the
fold-and-thrust directly east of the Sapphire and Lombard slabs referred to as
the Helena Salient. The ISr = 0.706 line represents the western edge of
Precambrian basement (Modified from Sears and Hendrix, 2004).

Hendrix (2004) show that the LCL was reactivated as a major left-Lateral transpressional

shear zone from the Late Cretaceous to the Early Eocene (-88-55 Ma). In their model,

left-Lateral motion along the LCL accommodated clock-wise rotation and eastward

displacement of the LEH, Sapphire, and Lombard thrust slabs in western Montana from









Late Cretaceous to Early Eocene time. In addition, Sears and Hendrix (2004) attribute

development of the Helena and Alberta (southern Alberta and northwestern Montana)

thrust salients to clock-wise rotation of these major thrust slabs along the LCL during this

time.

Late Cretaceous Magmatism

Voluminous magmatism was broadly synchronous with crustal shortening and

thickening in eastern Idaho and western Montana. In eastern Idaho and westernmost

Montana the Bitterroot lobe (-14,000 km2) of the Idaho-Bitterroot Batholith was

emplaced into the accreted Seven Devils-Wallow terrane and Belt Supergroup largely

from -120-52 Ma (see Fig. 1-1, Bickford et al., 1981; Hyndman, 1984; Foster and

Fanning, 1997; Foster et al., 2001). Early (-120-70 Ma) "deep level" quartz diorite and

tonalite plutons of the western Bitterroot lobe were emplaced at crustal depths up to -25

km (Hyndman et al., 1988; Foster et al., 2001). In the central and eastern Bitterroot lobe

"main phase" granodiorites and two-mica granites were emplaced at progressively

shallower crustal depths to the east (-25-10 km) between 65-52 Ma. In addition, several

younger alkali-feldspar granite plutons were emplaced in the older batholith intrusions

from -50-46 Ma at very shallow crustal depths ranging from -7-1.5 km (Foster et al.,

2001).

East of the Bitterroot lobe of the Idaho-Batholith Batholith several smaller

batholiths, stocks, and plutons were emplaced into the Sapphire and Lombard thrust slabs

from -80-65 Ma at relatively shallow crustal depths (-1-15 km, Hyndman et al., 1988;

Kalakay et al., 2001). Major intrusions in this region include the Sapphire Batholith, the

Chief Joseph Batholith, the Flint Creek plutons, and the Boulder Batholith (Fig. 2-4).









The Boulder Batholith, the largest of these intrusions (-4,000 km2), intruded its

own volcanic cover (the Elkhom Mountains Volcanics) indicating a very shallow

emplacement depth, probably on the order of-1-10 km (Tilling et al., 1968). Individual

plutons and stocks within these large intrusions often exhibit a close spatial and temporal

relationship with thrust faulting in the Sapphire and Lombard thrust slabs. In several

places within the Flint Creek and Anaconda-Pintlar Ranges (i.e., the Sapphire thrust slab)

Late-Cretaceous plutons were preferentially emplaced along thrust faults ramps

(Hyndman et al., 1975; Kalakay et al., 2001). Older plutons commonly exhibit strong

solid-state deformation while slightly younger plutons emplaced along the thrusts (some

cross-cutting the thrusts) are undeformed; this relationship indicates the older plutons

were probably emplaced during thrusting while emplacement of the younger plutons

post-dated the thrusting (e.g., Hyndman et al. 1975; Hawley, 1975; Wallace et al., 1992;

Kalakay et al., 2001). Kalakay et al. (2001) reports similar relationships between Late-

Cretaceous plutons and thrust faulting in eastern Pioneer Batholith along the eastern front

of the Grasshopper thrust slab located directly south of the Sapphire thrust slab in

southwestern Montana. Similarly, the Boulder Batholith was emplaced into a major

thrust fault ramp system within the Lombard thrust slab in the Helena thrust salient at

-80-75 Ma (Fig 4., based on SHRIMP U-Pb zircon crystallization ages, Lund et al.,

2002). Lageson et al. (2001) suggest that emplacement of the Boulder Batholith (and the

Elkhorn Mountains Volcanics) at this time created a supercritical taper geometry in the

Lombard thrust slab facilitating further east-directed displacement of the slab.








































B. SAppti* bathoht 18. Hln vani fIlid H W. S
SPhIpsbwg batholMih '
10. Mainl Powel bathollh 2.y-l ^ ASSNOPP E
13 HuB! PLATE
15 30 M, lies3 ( p d ana
a 1j Otherg. 1981)
15 30 45 Kilom ret er .13

FOLD AND
do 4P 0 THRUST ZONE


Figure 2-4. Geologic map of western Montana, south of the Lewis and Clark line. Emphasis is on the location and names of Late
Cretaceous through middle Eocene granitic batholiths, stocks, plutons, and volcanic fields in the Sapphire thrust plate and
the frontal imbricate thrust zone which lies to the east (from Wallace et al., 1992).









High Grade Metamorphism during the Cretaceous

Grover et al. (1992) and House et al. (1997) document three major metamorphic

events that affected the areas adjacent to the Bitterroot lobe of the Idaho-Bitterroot

Batholith during the Late Early to Late Cretaceous. (1) Regional prograde

metamorphism at -100-80 Ma was synchronous with crustal shortening and partly

coincident with emplacement of deep level plutons in the western Bitterroot lobe (-120-

70 Ma, Foster et al., 2001). Pressure and temperature conditions during this

metamorphism have been estimated (based on phase assemblages and quantitative

thermobarometry) to be at Middle-amphibolite facies conditions, -5-6 kbar and -500-

6000C (House et al., 1997 and references therein). (2) Upper-amphibolite facies

metamorphism occurred synchronous to emplacement of main phase plutons in the

Bitterroot lobe at -64-53 Ma. Metamorphic grade decreases from upper-amphibolite

facies conditions (peak metamorphic conditions at -6-8 kbar and -650-7500C) directly

adjacent to the batholith to lower greenschist facies conditions at a distance of -30 km

(House et al., 1997; Foster et al., 2001). Local anatexis accompanied the upper-

amphibolite facies metamorphism near the batholith. Anatexis of Belt-equivalent quartz-

two feldspar gneisses was facilitated by the breakdown (dehydration) of muscovite within

underlying semi-pelitic schist under sillimanite-zone conditions (Foster et al., 2001). (3)

Later relatively low pressure (-4-6 kbar) high temperature upper-amphibolite facies

metamorphism and anatexis during isothermal decompressional accompanied

exhumation of the Bitterroot metamorphic core complex at -53-48 Ma (see below, House

et al., 1997; Foster et al., 2001).

Quantitative thermobarometric data were not available from areas east of the

Bitterroot-Idaho Batholith prior to this study. Some workers provide broad pressure-









temperature estimates based on the metamorphic phase assemblages of metasedimentary

strata adjacent to plutons. These estimates indicate localized low pressure Middle to

upper-amphibolite facies metamorphism at -1-4 kbar and -500-700C accompanied

emplacement of the plutons during the Late Cretaceous (-80-70 Ma, Hyndman, 1988, see

previous pressure-temperature constraints summarized in a Later section).

Eocene History

An episode of large-scale extension began in the Early to Middle Eocene with

inception of numerous metamorphic core complexes along the previously thickened

Sevier hinterland, north of the Snake River Plain (SRP) in present-day southern British

Columbia, northeastern Washington, northern Idaho, and western Montana (Foster et al.,

2001, Foster et al., 2006a). The large-scale core-complex related extension immediately

followed the end of crustal shortening (ca. -55 Ma) in the fold-and-thrust belt located

directly to the east of the hinterland (Fig. 1-2, Constenius, 1996; Foster et al., 2001; Sears

and Hendrix, 2004). Large-scale extension in these metamorphic core complexes was

facilitated by displacement along low-angle brittle/ductile normal-sense detachment fault

zones comprised of a brittle upper crustal portion rooted to a mid-crustal ductile portion;

the mid-crustal portion of these detachments are characterized by mylonitic shear zones

(Coney, 1980; Lister and Davis, 1989; Wemicke, 1992; Foster et al., 2001).

Early extension in the metamorphic core complexes of southern British Columbia

and the northwestern United States was accompanied by the voluminous Kamloops-

Colville-Challis-Absaroka (KCCA) magmatism in the back-arc region, -500-1000 km

inboard of the convergent plate margin (Armstrong and Ward, 1991; Morris et al., 2000;

Breitsprecher et al., 2003). The (KCCA) magmatism is expressed by a belt of calc-

alkaline and alkaline volcanic rocks extending from British Columbia into northwestern









Wyoming and roughly overlaps with areas affected by extension in the metamorphic core

complexes. The exhumed lower plates of many of the Eocene metamorphic core

complexes in southern British Columbia and the northwestern United States are

characterized by voluminous granitic plutons the same age as the (KCCA) volcanic rocks

(Armstrong and Ward, 1991; Foster et al., 2006a).

Regional Transtension and Large-scale Crustal Extension along the Lewis and
Clark Line

Large-scale crustal extension in the metamorphic core complexes of northeastern

Washington, northern Idaho, and western Montana was linked kinematically to regional

dextral transtensional along the Lewis and Clark line during the Eocene (LCL, Doughty

and Sheriff, 1992; Yin and Oertel, 1995; Foster et al., 2006a). The LCL is comprised of a

northwest-west trending -40-80 km wide zone of steeply-dipping strike-slip, oblique-

slip, and dip-slip faults that stretches from northeastern Washington into west-central

Montana (a distance greater than 800 km). Several major faults along the LCL have

accommodated repeated displacement during tectonic reactivations since inception of the

line during the Mesoproterozoic or earlier (Wallace et al., 1990; Sears and Hendrix,

2004). Regional dextral transtension and development of metamorphic core complexes

along the LCL began in Early to Middle Eocene time (at. -54-52 Ma) after a long-lived

episode of sinistral transpressional during the Late Cretaceous to early Eocene (ending at

-55 Ma) associated with east-directed crustal shortening in the fold-and-thrust-belt of

western Montana (Foster et al., 2001; Sears and Hendrix, 2004; Foster et al., 2006a). The

detachment fault zone of the metamorphic core complexes along the LCL are structurally

linked to one another via strike-slip and/or oblique-slip splays faults of the LCL; these









splay faults exhibit major Eocene aged dextral offset (Fig. 1-2, Hendrix and Sears, 2004;

Foster et al., 2006a).

Although the timing of large-scale Eocene extension in metamorphic core

complexes along the LCL is very similar, there are some significant differences in the

kinematics of extension in these metamorphic core complexes north and south of the

LCL. To the north, metamorphic core complexes were exhumed by paired (or

symmetric) east and west dipping detachment fault zone by ENE-WSW directed tectonic

unroofing (Fig. 1-2; e.g., Okanagan, Kettle, and the Priest River metamorphic

complexes). To the south, the BCC was exhumed by a single (or asymmetric) east

dipping detachment fault zone by ESE directed unroofing (Foster et al., 2006a). It has

been proposed that the ACC was exhumed along east-dipping detachment fault zone by

ESE directed unroofing as well (Kalakay et al., 2003; O'Neill et al., 2004, Foster et al.,

2006a). The Clearwater metamorphic core complexes is located within the LCL and

occupies a relay between two major strike-slip faults which link the Priest River

metamorphic core complex and the BCC (Fig. 1-2). The Clearwater metamorphic core

complexes was exhumed along single east-dipping detachment fault zone by top-to-the-

east directed unroofing, similar to the BCC (Foster et al., 2006a).














CHAPTER 3
THE ANACONDA METAMORPHIC CORE COMPLEX

Structural-metamorphic Domains of the Anaconda Metamorphic Core Complex:

The Anaconda metamorphic core complex (ACC) is located within the collapsed

Sevier hinterland province in western Montana, south of the Lewis and Clark line, east of

the Bitterroot metamorphic core complex, and west of the Boulder Batholith, and Helena

salient of the fold-and-thrust belt (Fig. 1-1 and 1-2). As in other Cordilleran

metamorphic core complexes (e.g., Coney, 1980), the ACC is subdivided into three

structural-metamorphic domains: (1) the high grade metamorphic and plutonic lower

plate footballl), (2) the unmetamorphosed or low-grade metamorphic, brittle faulted

upper plate (hanging wall), and (3) the detachment fault zone which juxtaposes and

separates the upper and lower plates.

In this chapter general descriptions are provided for each of the three structural-

metamorphic domains of the ACC. Detailed descriptions are provided for the Lake of the

Isle shear zone, a km-scale upper-amphibolite facies ductile shear zone mapped in the

northeastern Anaconda-Pintlar Range during this study. Detailed descriptions of the

individually mapped ACC lower plate units and a new 1:24,000 scale geologic map of

the study area are provided in Appendix B and D, respectively.

The Lower Plate

Rocks of the ACC lower plate are partly exposed in the cores of two broad

antiformal flexures, or domes, in the Flint Creek and northern Anaconda-Pintlar Ranges

that lie structurally beneath the detachment. These two antiforms are separated by a









synformal trough west of Anaconda, MT where upper plate rocks partly cover the lower

plate rocks (Fig. 3-1, O'Neill et al., 2004). Lower plate rocks in the Flint Creek and

Anaconda-Pintlar Ranges are comprised partly of metamorphic equivalents to the

Mesoproterozoic Belt Supergroup, Middle Cambrian, and in some areas, to Devonian to

Lower Cretaceous sedimentary strata. All of these metasedimentary strata were intruded

by multiple generations of Late Cretaceous to early-middle Eocene batholiths, plutons,

stocks, dikes, and sills (Fig. 3-1 and 3-2, Emmons and Calkins, 1913; Csejtey, 1963;

Stuart, 1966; Desmarais, 1983; Heise, 1983; Wallace et al., 1992; Lonn et al., 2003). A

new 1:24000 scale geologic map was produced from the ACC lower plate located in the

field area of this study. The map and a detailed description of the mapped lower plate

units are included in Appendix F and B, respectively.

In the Flint Creek Range, the intrusive rocks comprise epizonal biotite-homblende

granodiorite and granitic plutons of the Late Cretaceous Mount Powell and Philipsburg

Batholiths and the Royal Stock which intrude mostly Cambrian and Lower Cretaceous-

equivalent metasedimentary strata and Belt-equivalent metasedimentary strata in a few

areas (Fig. 3-1, Allen, 1966; Stuart, 1966; Hyndman et al., 1972; Lonn et al., 2003). In

the Anaconda-Pintlar Range, both Late Cretaceous and early to middle Eocene intrusive

rocks intrude mostly Belt and middle Cambrian-equivalent metasedimentary strata

(Desmarais, 1983; Wallace, et al., 1992; Lonn et al., 2003). The early to middle Eocene

intrusive rocks are much more voluminous and form a southward widening belt of biotite

granodiorite, biotite granite, and two-mica granite batholiths, stocks, and plutons along

the entire length of the Anaconda-Pintlar Range (Fig. 3-2, Desmarais, 1983; Wallace et

al., 1992).



































& volc c r s Pioneerk
Sj Range.







0 exploration well
red is exposed in the Anaconda-Pintlar eastern Flint Creek Ranges. The
Eocene greenschist Cretaceous Faults detachment
S C upes mylonite canic rocks
Tetary sediments Cretaceous n n n mal






SvoPhilipsburg Batholith, MPb = Mt. Powell Batholith, SLP = Storm Lake Stock
o Tertiaryt, Proterozoic. Palezoc
intrusive rocks PfJ Mesozo.: eaiimeriary \ other
rocks
0 exploration well
Figure 3-1. Geological Map of the Anaconda metamorphic core complex (ACC) in
western Montana. The exhumed ACC lower plate footballl) highlighted in
red is exposed in the Anaconda-Pintlar eastern Flint Creek Ranges. The
detachment zone bounds the eastern side of the lower plate and consists of the
Eocene greenschist facies mylonite and overprinting brittle detachment. The
ACC upper plate (hanging wall) lies mostly east of the lower plate and
detachment zone with limited exposes just west of Anaconda, MT between the
Anaconda-Pintlar and Flint Creek Ranges. RS = Royal Stock, Pb =
Philipsburg Batholith, MPb = Mt. Powell Batholith, SLP = Storm Lake Stock
pluton, SP = Sapphire Batholith, CJ = Chief Joseph Batholith, PB = Pioneer
Batholith, BB = Boulder Batholith, LCV = Lowland Creek Volcanic Field,
EHV = Elkhorn Mountains Volcanic Field. The current study area is marked
by the dashed box (Modified from Foster et al., 2006a).















Smigmatitic lower Bett


Deerlodge
valley


Sattenuated mafic sills

r sheeted Tertiary granodiorite

seismic line BL2 of A'
Vjiemelek R. Smithsnn 19 OO


w- --471- 1", Amoco Amoo .----..... E
". AMoA J acobson 1



5km BOULDER BATHOLITH 55km
complex of meta-Belt & plutons .
underlain by Paleoproterazoic crust -

15 km
Figure 3-2. Geologic cross section of the Anaconda metamorphic core complex, western Montana. The Eocene mylonite and brittle
detachment dip gently to the east beneath the Deerlodge Valley and project beneath the Boulder Batholith as indicated by
Amoco exploration wells and seismic reflection data (see text for details). Equal-area stereonet projection shows the
orientation of mylonite stretching lineations (solid squares) and fault slickenlines (open circles) from the detachment zone
(Stereonet data and cross section from courtesy of Kalakay).









Major structures exposed with the ACC lower plate include numerous folds and

low-angle thrust faults related to crustal shortening along the eastern edge of the Sapphire

(Skalkaho) thrust plate during the Late Cretaceous (Hyndman, 1975; Hyndman et al.,

1988; Wallace et al., 1992; Sears and Hendrix, 2004). The Georgetown thrust (Emmons

and Calkins, 1913), the largest of the thrust fault structures, places metamorphosed

Helena Formation (middle Belt Carbonate-equivalent) over Devonian, Pennsylvanian,

and Mississippian metasedimentary strata (>7000 m of stratigraphic offset) and marks the

western margin of the ACC lower plate from central Anaconda-Pintlar Range to the

northern Flint Creek Range (Fig. 3-2, Wallace et al., 1992; O'Neill et al., 2004). The

northern end of this thrust is cut by the Philipsburg Batholith in the western Flint Creek

Range which yielded concordant hornblende and biotite K-Ar cooling ages of -77-72 Ma

(Hyndman et al., 1972). Several other thrust faults and related folds involving Belt-

equivalent through Cretaceous metasedimentary strata have been documented in

throughout the Flint Creek and Anaconda-Pintlar Ranges (e.g., Emmons and Calkins,

1913; McGill, 1959; Flood, 1974; Heise, 1983; Baken, 1984; Wallace et al., 1992; Lonn

et al., 2003). A common characteristic in both the Anaconda-Pintlar and Flint Creek

Ranges is the close spatial and temporal relationship between thrusting and emplacement

of Late Cretaceous plutons (e.g., the Philipsburg Batholith and Georgetown thrust,

Emmons and Calkins, 1913; Wallace et al., 1992).

Other structures found within the ACC lower include high-angle, sometimes listric-

shaped, normal faults associated with the brittle detachment found mostly along the

eastern flanks of the Flint Creek and Anaconda-Pintlar Ranges (see below). The largest

normal fault exposed in the ACC lower plate is the Hidden Lake-Dry Creek fault zone









(HLCDFZ), a large east-dipping, listric-shaped, locally ductile normal fault zone that

strikes parallel to and lies directly east of the Georgetown thrust (Fig. 3-1, 3-2). The

HLDCF shows >450 m of offset in some areas and may have been active in the Late-

Cretaceous during intrusion of the Dora Thorn pluton of the Philipsburg Batholith

(Buckley, 1990) and during the Eocene, as listric splays of the HLDCF form ductile shear

zones in the middle Eocene Pintlar Creek Batholith in the northern Anaconda-Pintlar

Range (Wallace et al., 1992). The shallow portion of these splay faults may have been

part of the original east-dipping detachment breakaway zone or an east-dipping brittle

normal fault synthetic to the basal detachment (O'Neill et al., 2004).

The Mesoproterozoic Belt through Cretaceous-equivalent metasedimentary strata

exposed in the ACC lower plate have been subjected to at least two major episodes of

high grade metamorphism: (1) pervasive regional metamorphism that predated the

emplacement of voluminous epizonal Late Cretaceous to Middle Eocene intrusions and

(2) high temperature, lower pressure contact metamorphism associated with the

emplacement of the Late Cretaceous to Middle Eocene (Emmons and Calkins, 1913;

Csejtey, 1963; Stuart, 1966; Desmarais, 1983; Heise, 1983; Hyndman et al., 1988;

Wallace et al., 1992; Kalakay et al., 2003; Grice et al., 2005). In the Flint Creek Range, a

pervasive regional upper-amphibolite facies cordierite-bearing assemblage, found mostly

within pelitic Cretaceous-equivalent phyllite, is overprinted by andalusite-bearing

assemblages in contact aureoles surrounding the Late Cretaceous intrusions (Stuart, 1966;

Buck, 1990). The earlier regional metamorphic assemblages show an overall increase in

metamorphic grade from greenschist to upper-amphibolite facies from west to east across

the Flint Creek Range (Stuart, 1966). In the Anaconda-Pintlar Range, upper-amphibolite









facies sillimanite-cordierite bearing assemblages are common in Lower Belt-equivalent

pelitic metasedimentary strata (Greyson Fm.) surrounding Late Cretaceous granodiorite

and quartz diorite plutons and sills. These metasedimentary strata are locally migmatitic

and characterized by granitic leucosome, abundant sillimanite, cordierite, garnet, K-

feldspar, biotite, and no primary muscovite (Desmarais, 1983; Grice et al., 2005). This

phase assemblage overprints an earlier higher pressure kyanite-bearing assemblage in the

current study area in the northeastern Anaconda Pintlar Range (see description of the

LISZ below, Kalakay et al., 2003; Grice et al., 2005). In addition, andalusite-bearing

assemblages overprint cordierite-sillimanite-bearing assemblages in some pelitic

metasedimentary strata adjacent to early to middle Eocene intrusions in the northern

Anaconda Pintlar Range (Emmons and Calkins, 1913). Andalusite assemblages were not

documented in the contact aureoles of Late Cretaceous intrusions in the current study

area. As in the Flint Creek Range, the regional metamorphic grade shows a general

increase from middle to uppermost amphibolite facies from west to east across the

Anaconda-Pintlar Range (Kalakay et al., 2003). A summary of previous pressure-

temperature estimates from the lower plate, along with new thermobarometry from LISZ

in this study, are provided in chapter 5.

The Detachment Fault Zone

The high grade metamorphic and plutonic rocks of the lower plate and

unmetamorphosed upper plate rocks of the ACC are juxtaposed and separated by an east-

dipping low-angle brittle detachment that exhibits significant top-to-the-east-southeast

displacement (Fig. 3-1 and 3-2, O'Neill et al., 2002; Kalakay et al., 2003; O'Neill et al.,

2004). The detachment was first recognized and mapped by Emmons and Calkins (1913)

who noted its striking similarities with the "great Bitterroot fault," which was later









defined to be the low-angle detachment that bounds the Bitterroot metamorphic core

complex to the west (Hyndman, 1980). The detachment now has a documented strike

length of >100 km stretching from the northern flank of the Flint Creek Range south to

the Bighole Valley along the eastern flanks of the southern Anaconda-Pintlar Range

(Kalakay et al., 2003; O'Neill et al., 2004). In the north, the detachment terminates into

steeply-dipping strike-slip and oblique-slip splay faults of the Ninemile fault, part of the

greater Lewis and Clark line (Foster et al., 2006a). The southern termination of the

brittle-ductile detachment has not been well established. Along the eastern flanks of the

Flint Creek and Anaconda-Pintlar Ranges the detachment dips gently (-10-30) beneath

the Deerlodge Valley; the gentle dip of the detachment is collaborated by industry

exploration wells that intersected greenschist mylonites at the base of the Tertiary basin

fill in the western Deerlodge Valley at depths of <5 km (Fig 3-2, McLeod, 1987). In

addition, the downward projection of the detachment aligns well with sub-horizontal

seismic reflectors beneath the Boulder Batholith, suggesting the detachment flattens with

depth and continues to the east (Fig. 3-2, Vejmelek and Smithson, 1995; Foster et al.,

2006a). The detachment is not well exposed along the western margin of the ACC;

however, it is inferred in several places by the juxtaposition of brittlely faulted upper

plate rocks with ductily deformed metamorphic rocks and plutonic rocks. The western

part of the detachment probably originated as a series of east-dipping listric-shaped

normal faults east of a breakaway zone, which is inferred to be directly east of the

Georgetown thrust (O'Neill et al., 2004; Foster et al., 2006a).

In the northeastern Anaconda-Pintlar Range, within the current study area, the

brittle detachment overprints a -300-500 m thick greenschist facies mylonitic shear zone









comprised mostly of stretched two-mica granite, biotite granite, and granodiorite and

minor micaeous quartzite (Fig. 3-3, Emmons and Calkins, 1913; Kalakay et al., 2003;

Appendix D). Strain is heterogeneous in the granitoid mylonites and distributed into

numerous 1-2 m thick ultramylonite zones separated by 5-15 m thick zones of mylonite

and protomylonite (Foster et al., 2006a). The metamorphic grade of the granitoid

mylonites is indicated by brittle fractured feldspar porphyroclasts surrounded by a matrix

of plastically deformed quartz (Kalakay et al., 2003). In addition, micaeous quartzite

mylonites exhibit unannealed quartz grains with well-developed undulatory extension in

thin section (Fig. 3-4a); these features are indicative of lower to middle greenschist facies

metamorphism at temperatures < 400-4500C (Wells et al., 2000). In addition, the

greenschist mylonites contain kinematic indicators and shallow plunging mineral

stretching lineations that show top-to-the-east-southeast sense of movement; mineral

lineations in both the granitoid and micaeous quartzite mylonites are comprised of

stretched quartz ribbons (Fig 3-4a and 4b, 102-108, Kalakay et al., 2003; O'Neill et al.,

2004).

The greenschist facies mylonites exposed in the Mill and Clear Creek drainages are

cut by series of closely spaced (-0.1-1 km) east-dipping listric-shaped normal faults that

commonly become sub-horizontal with depth and tangential to the highly strained

ultramylonite zones in the mylonites (Kalakay et al., 2003). In some places, the deeper

portions of these brittle faults terminate into the detachment (O'Neill et al., 2004). Good

examples of these faults are found on the western walls of the Clear Creek drainage and

along the continental divide separating the Mill and Tenmile Lake drainages (Fig. 7-5,

Appendix D). Slickenline striations on these fault surfaces indicate top-to-the-east-









southeast (100-110) slip, consistent with motion in the greenschist mylonites (Kalakay

et al., 2003; Foster et al., 2006a).






















Figure 3-3. Photomosaic of the Anaconda metamorphic core complex. Photo directed to
the northwest showing the brittle detachment (dashed and barbed line) which juxtaposes
unmetamorphosed Tertiary upper plate rocks with the Eocene greenschist faces mylonite
zone and high grade lower plate rocks.

To the north, along the eastern flank of the Flint Creek Range, Eocene mylonites

formed in the eastern parts of the Mount Powell Batholith and Royal Stock, both

comprised mostly of Late Cretaceous granodiorite and granite plutons (Allen, 1966; Lonn

et al., 2003; O'Neill et al., 2004). These mylonites are cut by high-angle normal faults

similar to those exposed in the Clear and Mill Creek drainages in the northeastern

Anaconda-Pintlar Range (O'Neill et al., 2004). To the south, along the eastern flanks of

the central and southern Anaconda-Pintlar Range, Wallace et al. (1992) mapped and

described low grade mylonites in two mica granites and granodiorites of the Middle

Eocene Pintlar Creek Batholith. These mylonites are cut by a series of northeast-trending









steeply-dipping normal faults related to the detachment that juxtapose the mylonites with


unmetamorphosed Tertiary fluvial deposits.


Figure 3-4. Greenschist facies mylonites. A) Outcrop photo of stretched granodiorite
with greenschist facies mylonitic fabric, upper Clear Creek drainage. Pencil is
oriented parallel to the stretching lineation (T/P = 1000/100). B)
Photomicrograph of mylonitic micaeous quartzite, Short Peak, right is to the
ESE. Asymmetric mica fish with top-to-the-ESE sense of shear. Note
undulatory extinction in the large deformed quartz grains. Cross polar light;
field of view is -2 mm.

The Upper Plate

Structurally above the high grade metamorphic-plutonic lower plate, greenschist

facies mylonitic shear zone, and brittle detachment, the eastern ACC upper plate is

mostly composed of an array of asymmetrical fault-bound basins filled with


























I B I


Figure 3-5. Outcrop photos of late listric-shaped brittle normal faults in the greenschist
facies mylonite zone. A) Photo taken from the southeastern flank of Short
Peak looking towards the northwest showing a listric-shaped brittle fault that
juxtaposed mylonitic micaeous quartzite with biotite granodiorite. Symbols:
Dot within the circle indicates previous fault motion out-of-the-page and X
within the circle indicates previous fault motion "into-the-page." B) Photo
taken from eastern side of the upper Clear Creek drainage looking south
showing an east-dipping listric-shaped brittle fault cutting greenschist facies
two-mica granite mylonite (photo in B is courtesy of David Foster).

unmetamorphosed syn-extensional Tertiary sedimentary, volcaniclastic, and volcanic

rocks (Fig. 3-1, 3-2). The faults that bound the basins are listric-shaped and sometimes









sole into the low-angle detachment at depth. In the Deerlodge Valley, the

stratagraphically lowest rocks in the basins consists of moderately west tilted (-50-600),

poorly sorted, and poorly consolidated conglomerates, sandstones, breccias and mega-

breccias (Kalakay et al., 2003; O'Neill et al., 2004). These strata grade upwards into

progressively less tilted (-0-250) volcanic lava flows, volcanic tuffs, and volcaniclastic

units that are correlated with the early to middle Eocene Lowland Creek Volcanic (LCV)

sequence (-54-48 Ma, based on 40Ar/39Ar cooling ages from the volcanic units by

Isopolatov, 1997; Lewis, 1990; Kalakay et al., 2003). The significant upward decrease in

tilt of the basin fill strata indicates deposition during brittle extension of the ACC upper

plate in the early to middle Eocene (Kalakay et al., 2003). Clasts of greenschist facies

two-mica granite and granodiorite mylonite were found in stratagraphically high

volcaniclastic lahar deposits in the upper plate, directly south of the current study area.

However, the age or volcano-stratigraphic correlation of these volcaniclastic deposits is

not yet well established (Kalakay et al., 2003; Foster, per comm.).

Description of the Lake of the Isle Shear Zone

A sinuous, km-scale ductile shear zone was documented and mapped in the ACC

lower plate within the northeastern Anaconda-Pintlar Range during this study. The

ductile shear zone strikes approximately east-west and outcrops over a broad area from

immediately west of Storm Lake -15 km to the east near the base of Mount Haggin in the

upper Mill Creek drainage. In the Mount Haggin area, the easternmost part of the ductile

shear zone is truncated by granitoids and greenschist mylonites associated with the

Eocene detachment. The ductile shear zone has been named the Lake of the Isle shear

zone (LISZ) herein after a small lake in the central part of the shear zone (see Appendix

D).









Mesoproterozoic Belt Supergroup and Middle Cambrian-equivalent

metasedimentary in the LISZ have been subjected to pervasive ductile deformation under

middle to upper-amphibolite facies conditions and exhibit a well-developed transposed

metamorphic foliation throughout the structure; the dip of the foliation varies from gentle

to moderate in east and west and is sub-vertical in the central LISZ (Appendix F). As a

result, the metasedimentary strata deformed in the LISZ are strongly attenuated (ductily

thinned) and characterized by dramatically reduced original stratigraphic thicknesses,

common mesoscopic boudins, and recumbent, locally isoclinal mesoscopic folds. Late

Cretaceous and early to middle Eocene plutons, dikes, and sills intrude the LISZ along its

entire length. The spatial relationship between these intrusions and the LISZ is described

below along with several aspects of the ductile deformation and high grade

metamorphism that were briefly summarized here.

Description and Distribution of Metamorphic and Structural fabrics

Metamorphic foliations and gneissic banding

Metasedimentary strata deformed and metamorphosed in the LISZ are correlated

with metamorphosed Mesoproterozoic Belt Supergroup and Middle Cambrian

sedimentary strata (Appendix B). Belt Supergroup-equivalent metasedimentary strata

mapped in the LISZ include the metamorphosed Greyson Formation (Lower Belt),

Ravalli Group, Helena Formation (Middle Belt Carbonate), and Missoula Group. Middle

Cambrian-equivalent metasedimentary strata mapped include the metamorphosed

Flathead Formation, Silver Hill Formation, and Hasmark Formation.

All of the meta-Belt and meta-Cambrian metasedimentary strata in the LISZ exhibit

a metamorphic foliation and/or gneissic (compositional) banding. The pelitic schist

correlated with Greyson Formation (Lower Belt) shows a strongly developed foliation









defined by the parallel alignment of abundant coarse-grained biotite with less abundant

fine-grained sillimanite fibrolite. Some exposures of the meta-Greyson pelitic schist in

the southern study area have a foliation comprised of biotite + sillimanite fibrolite +

muscovite. In the central and eastern study area the pelitic schist grades to a pelitic

paragneiss (also correlated with the metamorphosed Greyson Fm.) characterized by a

distinct gneissic banding. The gneissic banding in these strata is defined by alternating

quartzite and pelitic layers, both typically ranging from -1-20 cm in thickness. The

quartzite layers of the gneissic banding are comprised of medium to coarse-grained light-

to-dark grey quartzite with minor biotite. The pelitic layers are medium to coarse-grained

and comprised largely of (in increasing relative abundance) cordierite + garnet +

plagioclase + biotite + K-feldspar + quartz + sillimanite fibrolite. Within the pelitic

layers very abundant sillimanite fibrolite and much less abundant biotite (note the lack of

muscovite) comprise a strongly developed foliation parallel to the gneissic banding (Fig.

7-6).

The biotite quartzite paragneiss (metamorphosed Ravalli Group) often exhibits a

moderately to strongly developed gneissic banding. This gneissic banding is defined by

alternating quartzite and more pelitic layers, both ranging in thickness from a few

centimeters to a few meters in some places. The quartzite layers are composed of light to

medium-grey medium-grained quartzite with minor biotite. The more pelitic layers are

typically medium-grained and largely comprised of (in increasing relative abundance) K-

feldspar + biotite (commonly chloritized) + plagioclase + quartz + muscovite. In both the

quartzite and more pelitic layers aligned biotite defines a weakly to moderately developed

foliation parallel to the gneissic banding (Fig. 7-7a). In addition, in some exposures the









foliation in the more pelitic layers is comprised of aligned biotite + sillimanite fibrolite +

muscovite.

The calc-silicate paragneisses correlated with the metamorphosed Helena

Formation and Missoula Group (middle and upper Belt, respectively) are also commonly

characterized by a prominent gneissic banding. Within these calc-silicate paragneisses,

gneissic banding is defined by alternating light-grey to light-green and dark-green layers.

The light-grey to light-green layers are mostly composed of quartz + calcite chlorite

while the dark-green layers are comprised of diopside + calcite + quartz + chlorite

tremolite sericite white mica. Light-grey to light-green and dark-green layers are both

fine-grained and typically range in thickness from <1 mm to -10 cm (Fig. 7-7b). Aligned

chlorite defines a moderately developed foliation in both of these layer types which is

oriented parallel to the gneissic banding. A thin section prepared from the dark green

layering in the Missoula Group-equivalent calc-silicate paragneiss exhibits highly altered

tremolite phenocrysts surrounded by a matrix rich in clinopyroxene (Fig. 8). Near the

base of both the Helena Formation and Missoula Group the calc-silicate paragneisses

grade into layers of alternating biotite schist and biotite-rich quartzite where biotite

chlorite define a moderate to strongly developed foliation parallel to gneissic banding.

The metamorphosed middle Cambrian-equivalent strata within the LISZ lack

gneissic banding but do exhibit variably developed metamorphic foliations. The white to

pink coarse-grained quartzite correlated with metamorphosed Flathead Formation shows

a weakly to moderately developed foliation in some areas defined solely by the parallel

alignment of muscovite. Fine-grained biotite schist correlated with metamorphosed









Silver Hill Formation exhibits a moderately developed foliation composed of aligned

biotite + sillimanite. The white coarse-grained marble correlated with the Hasmark


Figure 3-6. Outcrop photograph showing the strong gneissic banding of the
metamorphosed Greyson Formation in the LISZ. The inset shows a close up
of the outcrop exhibiting large garnet porphyroblasts encased by a
metamorphic foliation comprised of abundant sillimanite fibrolite with less
biotite. Photograph is directed to the east.





















---._.2---p.-,..
,~ -~C J~imT



*
i r .
L i7d iW L-IVI


Figure 3-7. Outcrop photographs of gneissic banding commonly observed in the LISZ.
A) Gneissic banding common in the metamorphosed Ravalli Group. Rock
hammer handle for scale. B) Gneissic banding common in the
metamorphosed Missoula Group. The pencil sharpener is for scale.

Formation lacks a visible foliation in most places observed. However, a few outcrops

exhibit a weak foliation defined by the parallel alignment of sparse chlorite.

In the central and eastern study area, a prominent deformed quartz diorite sill is

emplaced within Belt and Middle Cambrian-equivalent metasedimentary strata and









exhibits a moderately to strongly developed solid-state foliation. The solid-state foliation

is defined by the parallel alignment of elongate hornblende, K-feldspar, plagioclase, and



Cpx





Trem










Figure 3-8. Photomicrograph of the Missoula Group-equivalent calc-silicate paragneiss.
Note large altered tremolite grains surrounded by fine-grained clinopyroxene
in the matrix. Photomicrograph was taken in cross polar light. Trem =
Tremolite, cpx = Clinopyroxene. Field offiew is -3 mm.

quartz (Fig. 3-9). This foliation is concordant to (parallel with) the metamorphic

foliations and gneissic banding in the adjacent attenuated Belt and middle Cambrian-

equivalent metasedimentary strata.

Mesoscopic-scale folds

Mesoscopic-scale folds are found in both the meta-Belt and meta-Cambrian

metasedimentary strata in the LISZ, but especially in meta-Belt metasedimentary strata.

These mesoscopic folds typically have geometries that range from open (60-120

interlimb angles) to isoclinal (0-100 interlimb angle). The hinge lines of the folds are

oriented parallel to the strike of the metamorphic foliation (and/or gneissic banding) in

the metasedimentary strata and their plunges vary from steep to shallow. The orientation









of the axial plane with respect to the foliation ranges from upright (70-90) to recumbent

(0-10) (e.g., Van Der Pluijm and Marshak, 2004, p. 243-247). In addition, the style of

mesoscopic folding in the LISZ ranges from parallel (Class 1B, constant layer thickness

maintained) to Class 3 type folding (significant thickening in fold hinge region and

thinning in the limbs, see Ramsay, 1967 for fold classification). Figures 3-10 and 3-11

displays common mesoscopic fold geometries found in the Belt metasedimentary strata

of the LISZ.

It is important to note here that the overall geometry of the mesoscopic folds

markedly changes across the east-west striking LISZ (and the contact with the deformed

quartz diorite sill). Near the center of the LISZ, directly adjacent to the deformed quartz

diorite sill, mesoscopic folds are tight to isoclinal and almost exclusively recumbent.

This fold geometry is especially common in the Greyson pelitic paragneiss adjacent to

the sill where the folds so flattened and transposed it is often difficult to distinguish the

folds from the planar gneissic banding. Away from the deformed quartz diorite the

mesoscopic folds become progressively more open and are typically not recumbent.

Mesoscopic-scale boudins

Mesoscopic boudins are also common in the Belt-equivalent metasedimentary

strata within the LISZ. In the Greyson Formation and Ravalli Group-equivalent

metasedimentary strata these boudins are found in quartzite layers. In the calc-silicate

paragneisses correlated with the metamorphosed Helena Formation and Missoula Group

the boudins are found within the lighter-colored more quartz-rich layers. Boudins within

the Belt-equivalent metasedimentary strata are blocky to tablet or lozenge-shaped and

range in length and width from -5 cm to 1-1.5 m and -5-25 cm, respectively. In

addition, when asymmetrical the longer axes of the boudins are usually oriented parallel









sub-parallel to the metamorphic foliation (and/or gneissic banding) in the

metasedimentary strata.


Figure 3-9. Outcrop photograph of the deformed quartz diorite sill in the center of the
LISZ. The sill exhibits a strong solid-state foliation (sub-vertical in
photograph) concordant to metamorphic foliation in the adjacent
metasedimentary country rocks. This relationship indicates the sill was
emplaced into the LISZ prior to or during some phase of solid-state
deformation. The photo is directed to the west. Rock hammer handle for
scale.

As in the case of the mesoscopic folds, the overall geometries of the mesoscopic

boudins change in a direction perpendicular to the overall strike of the LISZ. In the









central LISZ, adjacent to the deformed quartz diorite sill, the boudins are more flattened,

lozenge-shaped, and are characterized by large aspect ratios (i.e., length to width ratio;

Fig. 3-12a). However, away from the central LISZ and deformed quartz diorite sill the

mesoscopic boudins are typically block-shaped and shorter, having smaller aspect ratios

(Fig. 3-12b).


Figure 3-10. Outcrop photographs of mesoscopic-scale folds found in Belt equivalent
metasedimentary strata deformed in the LISZ. A) Nearly isoclinal class 3
folds in meta-Greyson Formation paragneiss. B) Mesoscopic folds in the
meta-Ravalli Group biotite quartzite gneiss.









Shear sense

The sense of shear for the LISZ is shown by shear-sense indicators found mostly in

the meta-Greyson pelitic schist and paragneiss and the deformed quartz diorite sill in the

central part of the shear zone. In the meta-Greyson, large garnet and cordierite

porphyroblasts with pressure shadows "tails" form these shear-sense indicators.

Flattened plagioclase and K-feldspar porphyroclasts form the shear-sense indicators in

the deformed quart diorite sill. In most observations, the porphyroblasts of the meta-

Greyson and porphyroclasts of the deformed quartz diorite sill formed approximately

symmetric sigma-type shear-sense indicators with pressure shadow tails aligned with the

median line of the porphyroblast (the median bisects the porphyroblast) indicative of pure

(or flattening) shear (Fig. 3-13a, Passchier et al., 1990). However, in a few observations

outcrops located directly east and west of the Lake of the Isle show garnet porphyroblasts

that form asymmetric sigma-type shear indicators (with offset pressure shadow tails)

showing evidence for left-lateral simple shear (Fig. 3-13b).

Description and Distribution of Metamorphic Phase Assemblages and Textures

All Belt and Middle Cambrian-equivalent metasedimentary strata in the LISZ bear

middle to upper-amphibolite facies phase assemblages. The phase assemblages were

briefly summarized above in the descriptions of the metamorphic fabrics. Note that a

description of these phase assemblages can be found in Appendix B. Here, descriptions

and distributions of key or "index" metamorphic phase assemblages and important

textural features are summarized to facilitate a later discussion concerning the PT history

of the LISZ during the Late Cretaceous.















































Figure 3-11. Outcrop photographs of mesoscopic-scale folds found in Belt-equivalent
metasedimentary strata deformed in the LISZ. A) Mesoscopic folds in meta-
Helena Formation calc-silicate gneiss on the eastern Flank of Mount Tiny.
Pencil is for scale. B) Mesoscopic Z-fold found in the meta-Ravalli Group
biotite quartzite gneiss. Pencil is for scale.

Metamorphic phase assemblages and textures in the meta-Greyson Formation

Pelitic strata correlated with the Greyson Formation exhibit significant changes in

metamorphic phases and textures along a transect perpendicular to the strike of the LISZ









and to the contact with the deformed quartz diorite sill. These changes are summarized in

the Table 3-1 which includes brief descriptions of the metamorphic phase assemblages

and important textures for six samples collected from the metamorphosed Greyson.


Figure 3-12. Outcrop photographs of mesoscopic boudins in the LISZ. A) In the outer
parts of the shear zone, away from the contact with deformed quartz diorite
sill, mesoscopic boudins are commonly blocky. B) In the center of the LISZ,
near the contact with the deformed quartz diorite sill boudins are more flat and
ellipsoidal.









Samples listed in Table 3-1 are arranged in order of decreasing distance from the contact

with the deformed quartz diorite sill in the central LISZ. Sample localities are shown in

Figure 3-16.

In the southernmost part of the study area, away from the contact with quartz

diorite sill and the central LISZ, the pelitic Greyson-equivalent schist bears abundant

coarse-grained biotite, relatively small euhedral garnets (-1-5 mm in diameter), K-

feldspar, fine-grained sillimanite fibrolite, and primary muscovite (e.g., WG04-108,

Table 3-1). This phase assemblage and general textural features are maintained within

the pelitic Greyson schist north towards the central LISZ until -1 km from contact with

the deformed quartz diorite sill. Here, the pelitic schist grades to the pelitic paragneiss as

described above. At this distance from the sill, the pelitic paragneiss contains much less

abundant biotite, relatively large (-5-10 mm in diameter) inclusion-rich subhedral

garnets, K-feldspar, relatively coarser-grained sillimanite fibrolite, and no primary

muscovite (e.g., WG04-026, Table 3-1).

Further north, at a distance of -0.5 km from the contact with the deformed quartz

diorite sill, some significant changes occur in the phase assemblage and texture of the

pelitic paragneiss. Here, the pelitic paragneiss contains much less biotite, relatively large

inclusion-rich subhedral to anhedral garnets, K-feldspar, very abundant relatively coarse-

grained sillimanite fibrolite, large sigmoidal-shaped cordierite porphyroblasts, and no

primary muscovite (e.g., ME-231, Table 3-1). In addition, at this distance from the

deformed quartz diorite sill, and closer, the pelitic paragneiss is migmatitic and contains

granitic leucosome comprised of (in increasing relative abundant) K-feldspar, quartz, and

albite. This leucosome is commonly found in boudin necks within boudinaged quartzite













































Figure 3-13. Outcrop photographs of shear sense indicators in the meta-Greyson
Formation deformed in the LISZ. A) Symmetric porphyroblast shear sense
indicator showing pure (flattening) shear. Hand lens for scale, photograph
courtesy of Heather Bleick. B) Asymmetric sigma-type garnet porphyroblast
shear sense indicator showing left lateral simple shear sense. Pencil is for
scale. Note symmetric shear sense indicators such as shown in A were
observed much more frequently than left later sense of shear indicator as
shown in B.

horizons and as thin elongate pods within the metamorphic foliation and/or gneissic

banding of the pelitic paragneiss (Fig. 3-14). In thin section, the leucosome is found

within the pressure shadows of garnet and cordierite porphyroblasts and as thin elongate









pods or lenses within the sillimanite fibrolite-biotite foliation (Fig. 3-15a). In addition,

the leucosome is found in thin section as thin "fingers" or veinlets cross-cutting garnet

porphyroblasts extending into the sillimanite fibrolite-biotite foliation (see the detailed

description of ME-231 included in Appendix C).

Based on the information summarized above and several other field observations

three major mineral/textural zones have been mapped within the pelitic schist and

paragneiss correlated with the metamorphosed Greyson Formation. These three zones

are shown on a simplified geologic sketch map in Figure 3-16. Zone 1 corresponds to

pelitic schist that contains abundant coarse-grained biotite, fine-grained sillimanite

fibrolite, relatively small euhedral garnets, and primary muscovite. Zone 2 corresponds

to pelitic paragneiss containing, less abundant biotite, fine-grained sillimanite fibrolite,

larger inclusion-rich subhedral to anhedral garnets, and no primary muscovite. Zone 3

represents pelitic paragneiss that contain little biotite, relatively abundant coarse-grained

sillimanite fibrolite, large inclusion-rich subhedral to anhedral garnets, large cordierite

porphyroblasts, and no primary muscovite. K-feldspar is found in all three mineral-

textural zones within the meta-Greyson.

Relict and fresh kyanite in the upper Meta-Ravalli Group

The uppermost -100-200 m section of Ravalli Group-equivalent metasedimentary

strata mapped is more pelitic than the rest of the section. These more pelitic strata are

well exposed near a small unnamed lake at the head of the Twin Lakes Creek drainage in

the western LISZ, southwest of Storm Lake, just north of the continental divide

(Appendix D). Here, the more pelitic upper Ravalli metasedimentary strata contain

abundant white to light cream-colored pseudomorphs. In outcrop, these pseudomorphs

form elongate blade-shape cross-sections when cut parallel to their longer axes. These









elongated pseudomorph cross-section range in length and width from -1-4 cm and -2-5

mm, respectively. When cut perpendicular to their longer axes, the pseudomorphs form

diamond-shaped cross-sections typically ranging from -3-10 mm in the longer dimension

(Fig. 3-17a).


Figure 3-14. Outcrop photographs of meta-Greyson migmatitic paragneiss showing
granitic leucosome commonly found in pressure shadows between quartzite
boudins. Upper: Field notebook for scale; upper photograph is courtesy of
Heather Bleick. Lower: Rock hammer is for scale.






52






















Pressure shadow















Figure 3-15. Photomicrographs of thin sections of the meta-Greyson migmatitic
paragneiss exhibiting granitic leucosome (sample ME-231). A) The
leucosome cross-cuts garnet porphyroblast. B) The leucosome is also found in
the pressure shadows of large cordierite porphyroblasts. Note the similar
placement of the leucosome in pressure shadows at both the outcrop and thin
section scale. gt = garnet, crd = cordierite, sill = sillimanite.
Photomicrographs were taken in cross polar light.

In thin section, the pseudomorphs are almost entirely composed of fine-grained

sericite white mica (Fig. 3-17b). Relatively coarse-grained muscovite laths commonly

form corona structures around the pseudomorphs. Small relicts of the original mineral









remain within the pseudomorphs as moderately to highly altered, but high relief

fragments that are typically < 0.5 mm in diameter or longest dimension (Fig. 3-17b).

Based on the blade-shape form (when cut parallel to their longer axes), the high relief of

the internal relict mineral material, and muscovite corona structures the pseudomorphs

exposed at this locality likely replace kyanite. In addition, small (< 1 mm) blade-shaped

mineral grains are found scattered throughout the matrix between large kyanite

pseudomorphs (Fig. 3-18). The small blade-shaped grains found in the matrix between

the large pseudomorphs are also kyanite, but apparently less altered (Nesse, 1991).

Spatial Relationship Between the Lake of the Isle Shear Zone and Late Cretaceous
Intrusions

The spatial relationship between major Late Cretaceous intrusions and the LISZ is

described in some detail in Appendix B. Here, a brief summary is given for the important

spatial relationships. The LISZ was intruded by four major Late Cretaceous intrusions

(see map in Appendix F). In the west, the LISZ was intruded by the quartz diorite and

granodiorite plutons of the SLS. Both of these intrusions are undeformed and cross-cut

the high grade and deformed meta-Belt and meta-Cambrian metasedimentary strata of the

LISZ indicating a post-kinematic relationship to the shear zone. In the east, a quartz

diorite pluton (similar to the quartz diorite of the SLS) cross-cuts the Mill Creek nappe

indicating a post-kinematic relationship to the LISZ as well. Between these two areas, a

quartz diorite sill was emplaced into the deformed metasedimentary strata of the central

LISZ. The quartz diorite sill is highly deformed and exhibits a well-developed foliation

in many places that is concordant to the foliation in the adjacent metasedimentary country

rocks indicating a pre or syn-kinematic relationship to the shear zone (see Appendix B

and F).













Table 3-1. Description and distribution of key mineral phases and textures in the metamorphosed Greyson Formation.

Sample Distance from Key phases Mineral/textural Textures
deformed qd sill (km) zone (Fig. 3-16)
WG04- bt is very abundant and fresh; grt is euhedral, no
108 4.0 bt, grt, kfs, ms, sill 1 embayments; no cordierite; fine-grained sill fibrolite
WG04- possible kyanite relicts; fine-grained sill fibrolite; grt is
095 1.2 bt, grt, kfs, ms, sill 1 v. euhedral; v. abundant ms
WG04- ms present; little sill, but fine-grained variety; bt is v.
101b 1.2 bt, grt, kfs, ms, sill 1 abundant
little sill, fine-grained fibrolite; possible kyanite
pseudomorphs; very abundant ms; rock is coarse
WG04- grained; bt is most abundant; smaller garnets <3 mm
099 1.2 bt, grt, kfs, ms, sill 1 in diameter
WG04- grt is subhedral and inclusion rich; coarser sill
026 0.8 bt, grt, kfs, sill 2 fibrolite; lacks ms
bt less abundant and is altered; coarser sill fibrolite is
very abundant; larger grt and large crd; crd have
bt, grt, kfs, sill, crd, abundant fine-grained sill fibrolite inclusions; lacks
ME-231 0.4 leuco 3 ms; abundant leuco
Note: qd = quartz diorite, bt = biotite, grt = garnet, ms = muscovite, kfs = K-feldspar, sill = sillimanite, crd = cordierite, and leuco = granitic
leucosome.













113'15'


46'7.5'


Figure 3-16. Simplified geologic sketch map of the ACC lower plate exposed in the current study area showing the metamorphic
mineral/textural zones 1, 2, and 3 mapped in the meta-Greyson Formation adjacent to the deformed quartz diorite sill. The
meta-Greyson Formation is brown, the sill is green, and uncorrelated Lower Belt pelitic rocks are light tan. Note the
localities of samples discussed in the text and described in Table 3-1 (see text for details).


















































Figure 3-17. Kyanite pseudomorphs in the upper meta-Ravalli Group. A) Outcrop
photograph of kyanite pseudomorphs. This outcrop is located in the western
LISZ (see Fig. 3-16 and sample locality for sample WG04-113). Pen for is
scale. B) Photomicrograph from a thin section of meta-Ravalli Group sample
WG04-113 showing large kyanite pseudomorph made of fine-grained sericite
white mica surrounded by relatively coarse-grained muscovite corona
structure. Photomicrograph was taken in cross polar light. Field of view is -3
mm.


;c ;;:-A^^~Ij
"j r "^^
i*>, ^.-Jtafc Vt-Staf-a.-






57





.. ...- .-
I -







5Z '5,,.4i --' ae"-., .... :- :. _,





Figure 3-18. Photomicrograph showing fresh kyanite in the upper meta-Ravalli Group.
Thin section of WG04-112 shows fine-grained fresh kyanite scattered through
the matrix of between large kyanite pseudomorphs. Field of view is -3 mm.














CHAPTER 4
U-PB ZIRCON GEOCHRONOLOGY

Purpose and Strategy

U-Pb zircon geochronology was employed in this study to provide age constraints

on upper-amphibolite facies metamorphism, anatexis, and ductile deformation in the

Lake of the Isle shear zone (see Chapter 3). Three samples were chosen from the

geochronology. WG04-114 was collected from the undeformed granodiorite of the

Storm Lake Stock (SLS) in the western LISZ (Appendix F). Here, the undeformed

granodiorite obliquely crosscuts the deformed metasedimentary strata of the western

LISZ, indicating that emplacement of the intrusion postdated upper-amphibolite facies

metamorphism and ductile deformation in the LISZ. A U-Pb zircon crystallization age

from WG04-114 would provide a minimum age limit for upper-amphibolite

metamorphism and ductile deformation in the LISZ. Ug-1 was collected from a quartz

diorite sill in the central LISZ where the sill is emplaced within a sequence of upper-

amphibolite grade meta-Belt and meta-Cambrian metasedimentary strata. The quartz

diorite sill is deformed and exhibits a well-developed solid-state foliation concordant

with the foliation in the adjacent metasedimentary country rocks. Such a relationship

indicates (1) emplacement of the sill predated upper-amphibolite metamorphism and

ductile deformation in the LISZ or (2) the sill was emplaced synkinematic to these

events. A U-Pb zircon crystallization age from the deformed quartz diorite sill sample

Ug-1 would provide an upper age limit or direct age constraint for upper-amphibolite

facies metamorphism and ductile deformation in the LISZ. Finally, sample WG05-02









was also collected from granitic leucosome found in the migmatitic pelitic paragneiss

(metamorphosed Greyson Fm., Lower Belt) in the central LISZ. Here there is evidence

for anatexis during the upper-amphibolite facies metamorphism and ductile deformation

(see Chapter 3). A reliable U-Pb zircon crystallization age from the leucosome sample

WG05-02 would provide a direct age constraint on upper-amphibolite facies

metamorphism, anatexis, and ductile deformation of metasedimentary strata in the LISZ.

Relevant Previous U-Pb Zircon Geochronology

Prior to this study, there were no U-Pb geochronological data available to provide

age constraints for the upper-amphibolite facies metamorphism and ductile deformation

in the LISZ. However, Desmarais (1983) reported U-Pb geochronological data from two

deformed granodiorite intrusions in the southern Anaconda-Pintlar Range which exhibit

solid-state foliations concordant with foliations in metasedimentary country rocks. U-Pb

zircon 207Pb / 235U and 206Pb / 238U ages from these deformed intrusions were discordant

with upper and lower concordia intersects at 1780-1890 and 78-77 Ma, respectively.

Desmarais attributed the discordance nature of the zircons to inheritance from

Proterozoic cores and he interpreted the lower intersect ages to be minimum ages for

emplacement for the intrusions.

U-Pb Zircon Geochronology Results

Zircons were extracted from samples WG04-114, Ug-1 and, WG05-02 using

standard rock crushing, density, and magnetic separation techniques. Select zircons were

subsequently analyzed along with FC-1 standard zircons (in house 207Pb / 206Pb age =

1086.9 5.3, 207b / 235U age = 1091.5 13.4, and 206Pb / 238U age = 1096.7 21.7 Ma)

at the University of Florida using laser ablation multi-collector inductively coupled









plasma mass spectrometry (LA-MC-ICP-MS). Sample preparation, LA-MC-ICP-MS

analytical techniques, and data reduction are summarized in Appendix A.

A common lead correction was applied to the drift and fractionation corrected U-Pb

zircon data obtained from zircons of samples WG04-114, Ug-1, and WG05-02 to

evaluate the need for the correction (Appendix A). In the case of WG04-114 and Ug-1,

individual U-Pb ages calculated from uncorrected and corrected isotopic data were within

analytical error indicating negligible concentrations of common lead in these zircons; a

common lead correction was deemed unnecessary for these analyses. For WG05-02, a

common lead correction is necessary due to the antiquity of the zircons and because all of

these zircon gave discordant 207Pb / 235U and 206Pb / 238U ages.

WG04-114 (Storm Lake Stock Granodiorite)

Twenty-eight spot analyses were taken from twenty-two euhedral and inclusion

free zircons from sample WG04-114 (average size = 200-300 [tm), including five paired

core and rim analyses from a few larger zircons. Individual U-Pb ages for the WG04-114

zircons are summarized in Table 4-1. 207Pb / 206Pb, 207Pb / 235U, and 206Pb / 238U ages are

reported for the older zircons. 206Pb / 238U ages are given for younger zircons. Core

analyses are indicated by the subscript "c" in Table 4-1. All ages in Table 4-1 are

reported with 2a errors. Three zircon analyses from WG04-114 yielded anomalously old

ages (SLS-7c-c, SLS-8a, and SLS-17b-c). These zircons gave discordant 207Pb / 235U and

206Pb / 238U ages ranging from -1189-239 Ma and -599-88 Ma, respectively. However,

the 207Pb / 206Pb ages are more consistent and range from -2502-2243 Ma indicating a

Paleoproterozoic or Archean component to these zircons. The remaining twenty-five

analyses from WG04-114 gave fairly consistent Late Cretaceous 206Pb / 238U ages ranging

from -80-71 Ma and are interpreted to represent a single magmatic zircon population









related to emplacement of the SLS granodiorite (Table 4-1). Select zircons were (n= 22)

are shown plotted on a Tera-Wasserburg diagram and 206Pb / 238U weighted mean zircon

age plot in Figure 4-1. Sixteen of these WG04-114 zircons have similar 206Pb / 238U ages

and correspond to a 206Pb / 238U weighted mean age of 74.6 0.8 Ma (MSWD = 0.34,

20). This age is interpreted to be the age of emplacement for the SLS granodiorite.

Ug-1 (The Deformed Quartz Diorite Sill)

A total of twenty-four spot analyses were taken from twenty-four individual

euhedral and inclusion free Ug-1 zircons (average size = 200-300 [tm); core and rim

zircon analyses were not distinguished during the analyses. Individual U-Pb ages for the

Ug-1 zircons are summarized in Table 4-2. As the case of WG04-114, 207Pb / 206Pb,

207Pb / 235U, and 206Pb / 238U ages are reported for older zircons and only 206Pb / 238U ages

are given for younger zircons. All ages in Table 4-2 are reported with 2C errors. Two

Ug-1 zircons yield anomalously old ages (Ug-01_9 and Ug-01_ 1). These two zircon

analyses gave discordant 207Pb / 235U and 206Pb / 238U that range from -510-471 Ma and

-299-109 Ma, respectively. The remaining twenty-two Ug-1 zircon analyses yield Late

Cretaceous 206Pb / 238U ages of -77-71 Ma. These zircons are interpreted to be of a

single magmatic population which crystallized during the emplacement of the quartz

diorite sill. The twenty-two Late Cretaceous aged Ug-1 analyses are shown plotted on a

Tera-Wasserburg diagram and 206Pb / 238U weighted mean zircon age plot in figure 4-2.

Sixteen of these zircons correspond to a 206Pb / 238U weighted mean age of 75.0 + 0.8 Ma

(MSWD = 0.54, 2a). This age is interpreted to represent the age of emplacement for the

quartz diorite sill.













Table 4-1. U-Pb LA-MC-ICP-MS analytical results for WG04-114
Radiogenic ratios
Grain spot 207Pb 207Pb 206Pb
206Pb + 235U e 238U e.c.
SLS-1 0.0573 0.0025 0.0987 0.0066 0.0121 0.0005 0.08
SLS-2 0.0572 0.0009 0.1026 0.0051 0.0121 0.0005 0.11
SLS-3 0.0566 0.0006 0.0891 0.0045 0.0119 0.0005 0.12
SLS-4 0.0546 0.0008 0.0992 0.0062 0.0119 0.0005 0.09
SLS-5a 0.0572 0.0014 0.0930 0.0047 0.0116 0.0005 0.11
SLS-5b-c 0.0550 0.0017 0.0979 0.0054 0.0118 0.0005 0.10
SLS-6 0.0739 0.0032 0.1225 0.0071 0.0118 0.0005 0.07
SLS-7a 0.0614 0.0022 0.0975 0.0068 0.0118 0.0005 0.08
SLS-7b-c 0.0604 0.0021 0.1051 0.0072 0.0114 0.0005 0.07
SLS-7c-c 0.1648 0.0011 2.2235 0.0975 0.0975 0.0045 0.05
SLS-8a 0.1415 0.0064 0.2654 0.0206 0.0137 0.0006 0.03
SLS-8b-c 0.0877 0.0064 0.1434 0.0141 0.0117 0.0005 0.04
SLS-9-c 0.0554 0.0009 0.0809 0.0051 0.0115 0.0005 0.10
SLS-10 0.0517 0.0012 0.0904 0.0047 0.0119 0.0005 0.11
















Table 4-1. Continued.
Ages (Ma)
Grain spot 207Pb 207Pb 206Pb
206Pb 235U + 238U
SLS-1 77.4 3.4
SLS-2 77.8 3.5
SLS-3 76.5 3.3
SLS-4 76.3 3.4
SLS-5a 74.5 3.3
SLS-5b-c 75.4 3.3
SLS-6 75.8 3.3
SLS-7a 75.4 3.3
SLS-7b-c 73.1 3.2
SLS-7c-c 2502 11 1,188.5 30.2 599.9 26.4
SLS-8a 2243 76 239.0 16.4 88.0 3.9
SLS-8b-c 74.9 3.3
SLS-9-c 73.9 3.3
SLS-10 76.5 3.3













Table 4-1. Continued.
Radiogenic ratios
Grain spot 207Pb 207Pb 206Pb
206Pb + 235U e 238 e.c.
SLS-11 0.0562 0.0013 0.0944 0.0057 0.0114 0.0005 0.09
SLS-12 0.0514 0.0010 0.1056 0.0068 0.0117 0.0005 0.08
SLS-13 0.0515 0.0009 0.0955 0.0052 0.0114 0.0005 0.10
SLS-14a 0.0517 0.0010 0.0914 0.0058 0.0111 0.0005 0.08
SLS-14b 0.0519 0.0009 0.0973 0.0058 0.0124 0.0005 0.10
SLS-15 0.0535 0.0011 0.0820 0.0048 0.0115 0.0005 0.11
SLS-16 0.0798 0.0029 0.1478 0.0117 0.0116 0.0005 0.04
SLS-17a 0.0849 0.0063 0.1548 0.0164 0.0117 0.0005 0.03
SLS-17b-c 0.1502 0.0031 0.2996 0.0195 0.0140 0.0006 0.03
SLS-18 0.0593 0.0032 0.0970 0.0096 0.0117 0.0005 0.05
SLS-19 0.0598 0.0020 0.0970 0.0056 0.0117 0.0005 0.09
SLS-20 0.0515 0.0014 0.0880 0.0061 0.0117 0.0005 0.09
SLS-21 0.0540 0.0019 0.0931 0.0072 0.0120 0.0005 0.07
SLS-22 0.1104 0.0034 0.2049 0.0215 0.0123 0.0006 0.03














Table 4-1. Continued.


207Pb
206Pb


207Pb
_ 235U


2346 37 266.1


Ages (Ma)
206Pb
+ 238U
73.3
75.0
73.2
71.2
79.5
73.8
74.5
74.9
15.1 89.4
75.1
75.0
75.0
76.7
78.8


Grain spot

SLS-11
SLS-12
SLS-13
SLS-14a
SLS-14b
SLS-15
SLS-16
SLS-17a
SLS-17b-c
SLS-18
SLS-19
SLS-20
SLS-21
SLS-22


3.3
3.4
3.3
3.1
3.5
3.3
3.3
3.4
3.9
3.3
3.3
3.4
3.4
3.6


Note: e.c. = error correlation between 207Pb/235U and 206Pb/238U errors. Analyses in italics
were excluded from the weighted 206Pb/238U mean age calculation.













Table 4-2. U-Pb LA-MC-ICP-MS analytical results for Ug-1
Radiogenic ratios
Grain spot 207Pb 207Pb 206Pb
206Pb + 235U e 238U e.c.
Ug-O1_1 0.0429 0.0008 0.0683 0.0029 0.0116 0.0005 0.17
Ug-01_2 0.0331 0.0014 0.0527 0.0030 0.0114 0.0005 0.17
Ug-01_3 0.0445 0.0008 0.0679 0.0030 0.0110 0.0005 0.16
Ug-01_4 0.0506 0.0013 0.0771 0.0036 0.0112 0.0005 0.14
Ug-01_5 0.0564 0.0040 0.0889 0.0076 0.0113 0.0005 0.06
Ug-01_6 0.0529 0.0022 0.0824 0.0047 0.0114 0.0005 0.11
Ug-01_7 0.0482 0.0017 0.0786 0.0041 0.0118 0.0005 0.13
Ug-01_8 0.0590 0.0025 0.0952 0.0054 0.0118 0.0005 0.09
Ug-01_9 0.0912 0.0006 0.5902 0.0250 0.0475 0.0021 0.08
Ug-01_10 0.0616 0.0018 0.0998 0.0053 0.0114 0.0005 0.09
Ug-01_11 0.2825 0.0021 0.6538 0.0284 0.0170 0.0007 0.03
Ug-01_12 0.0490 0.0004 0.0802 0.0033 0.0116 0.0005 0.15
Ug-01_13 0.0678 0.0019 0.1132 0.0061 0.0119 0.0005 0.08
Ug-01_13 0.0789 0.0039 0.1335 0.0087 0.0120 0.0005 0.06
Ug-01_15 0.0565 0.0020 0.0952 0.0054 0.0118 0.0005 0.09
Ug-01_16 0.0560 0.0004 0.0914 0.0039 0.0116 0.0005 0.13
Ug-01_17 0.0947 0.0034 0.1602 0.0089 0.0116 0.0005 0.06
Ug-01_18 0.0754 0.0041 0.1336 0.0103 0.0116 0.0005 0.05
Ug-01_19 0.1107 0.0020 0.1912 0.0087 0.0120 0.0005 0.06
Ug-01_20 0.0607 0.0014 0.1003 0.0051 0.0117 0.0005 0.10
Ug-01_21 0.0951 0.0031 0.1610 0.0086 0.0119 0.0005 0.06
Ug-01_22 0.0561 0.0014 0.0865 0.0041 0.0111 0.0005 0.12
Ug-01_23 0.0679 0.0022 0.1055 0.0056 0.0113 0.0005 0.09
Ug-01_24 0.0798 0.0038 0.1329 0.0090 0.0117 0.0005 0.06














Table 4-2. Continued.


207Pb
206pb


207Pb
_ 235U


Grain spot

Ug-01_1
Ug-01_2
Ug-01_3
Ug-01_4
Ug-01_5
Ug-01_6
Ug-01_7
Ug-01_8
Ug-01_9
Ug-01_10
Ug-01_ 11
Ug-01_12
Ug-01_13
Ug-01_13
Ug-01_15
Ug-01_16
Ug-01_17
Ug-01_18
Ug-01_19
Ug-01_20
Ug-01_21
Ug-01_22
Ug-01_23
Ua-01 24


Ages (Ma)
206Pb
+ 238U
74.1
73.0
70.7
72.1
72.4
72.9
75.8
75.5
15.8 299.1
73.1
17.3 108.7
74.6
76.3
77.0
75.6
74.3
74.4
74.1
76.7
74.8
76.1
71.2
72.2
75.2


1449 13 471.0

3375 12 510.8


3.2
3.1
3.0
3.1
3.1
3.1
3.3
3.2
12.7
3.1
4.7
3.2
3.3
3.3
3.3
3.2
3.2
3.2
3.3
3.2
3.3
3.1
3.1
3.3


Note: e.c. = error correlation between 207Pb/235U and 206Pb/238U errors. Analyses in italics
were excluded from the weighted 206Pb/238U mean age calculation.














Table 4-3. U-Pb LA-MC-ICP-MS analytical results for WG05-02
Radiogenic ratios


Grain spot 207Pb
206Pb
L1 0.1027
L2 0.1703
L3 0.1761
L4 0.1308
L5 0.1446
L6 0.1057
L7 0.1367
L8 0.0988
L9 0.1466
L10 0.1708
L11 0.1003
L12 0.2232
L13 0.1074
L14 0.1042
L15 0.1086
L16 0.1870
L17 0.1032
L18 0.0983
L19 0.1597
L20 0.1038
L21 0.1634
L22 0.1016
L23 0.0950
L24 0.7816
L25 0.1014


2
2


0.0006
0.0012
0.0010
0.0008
0.0009
0.0006
0.0008
0.0007
0.0086
0.0010
0.0006
0.0019
0.0006
0.0008
0.0008
0.0010
0.0006
0.0006
0.0009
0.0006
0.0009
0.0006
0.0006
0.0043
0.0006


1


07Pb
'35Uj
1.2005 0.0601
1.0616 0.0678
6.9862 0.2908
1.8921 0.1317
5.4060 0.2574
3.1837 0.1949
1.4190 0.0740
0.3773 0.0384
0.0346
10.4532 0.4889
1.7837 0.0848
4.4270 0.2277
3.3357 0.1581
1.9734 0.1056
1.3191 0.0765
9.9195 0.4345
4.0275 0.1739
1.1089 0.0615
3.4682 0.1518
2.2842 0.1000
6.3418 0.2750
0.8736 0.0482
0.4689 0.0291
14.0487
0.9928 0.0487


206Pb
238U
0.1010
0.0588
0.3003
0.1465
0.2714
0.2197
0.0779
0.0498
0.0034
0.3830
0.1118
0.1176
0.1994
0.1215
0.0935
0.3066
0.2247
0.0772
0.1644
0.1636
0.2849
0.0758
0.0451
0.0508
0.0780


0.0045
0.0029
0.0131
0.0068
0.0123
0.0126
0.0035
0.0023
0.0007
0.0175
0.0050
0.0056
0.0088
0.0061
0.0042
0.0139
0.0098
0.0036
0.0073
0.0073
0.0129
0.0035
0.0022
0.0242
0.0036


e.c.
0.08
0.04
0.05
0.05
0.05
0.06
0.05
0.06
0.02
0.04
0.06
0.02
0.06
0.06
0.05
0.03
0.06
0.06
0.05
0.07
0.05
0.07
0.08
0.00
0.07














Table 4-3. Continued.


Grain spot

L1
L2
L3
L4
L5
L6
L7
L8
L9
LIO
L11
L12
L13
L14
L15
L16
L17
L18
L19
L20
L21
L22
L23
L24
L25


Ages (Ma)


207Pb
206U
1671
2559
2616
2105
2280
1723
2184
1601

2564
1628
3004
1754
1698
1776
2712
1680
1592
2450
1689
2488
1651
1525


1647 11 700.1


207Pb
235U
800.8
734.6
2109.7
1078.3
1885.8
1453.2
896.9
325.1

2475.8
1039.5
1717.4
1489.5
1106.5
854.1
2427.3
1639.8
757.6
1520.0
1207.4
2024.3
637.5
390.4


206Pb
238U
620.6
368.3
1693.0
881.1
1547.7
1280.3
483.9
313.4

2090.5
683.3
716.9
1172.3
739.0
576.4
1723.8
1306.8
479.5
981.3
976.7
1616.2
471.2
284.6


24.5 484.0


27.4
32.9
36.3
45.2
40.0
46.2
30.6
27.9

42.4
30.5
41.7
36.4
35.4
33.0
39.6
34.5
29.2
33.9
30.5
37.3
25.8
19.9


26.5
17.7
64.8
38.2
62.0
66.1
20.9
14.4

80.9
29.1
32.3
47.3
35.2
24.5
68.3
51.3
21.5
40.5
40.4
64.5
20.8
13.8

21.3


Note: e.c. = error correlation between 207Pb/235U and 206Pb/238U errors. Analyses in italics
were excluded from the age calculation shown in Figure 4-3b.













0.14


0.08


data-poM ens ellmse am 2 ugma


74 78 82 86 90 94
mUaiPb


Figure 4-1. Tera-Wasserburg plot and 206Pb / 238U weighted mean zircon age plot for
Storm Lake Stock granodiorite sample WG04-114. Only the zircon analyses
represented by the darkened ellipses in A and the darkened boxes in B were
used in the weighted mean 206Pb / 238U zircon age calculation. U-Pb isotopic
data were not corrected for common lead (see text).

WG05-02 (Leucosome from the Meta-Greyson Paragneiss)

A total of twenty-five spot analyses were taken from twenty-five individual

subhedral and inclusion free zircons of sample WG05-02; because of their small size

(average size <100 [m), exclusive rim analyses were not made from these zircons.


WG04-114, Storm Lake Stock granodiorite














86 82 7R .74


. 70













0.14 -


0.12


.7Pb D.10
mPb




D.06




mU/IPb

box heigts are 2
82
B

79

77

7 74


S72

69
20l6pb23eU weighted mean
67 zircon age = 75.0 0.8 Ma
(MSWD=O 0.54)
64


Figure 4-2. Tera-Wasserburg plot and 206Pb / 238U weighted mean zircon age plot for the
deformed quartz diorite sill sample Ug-1. Only the zircon analyses
represented by the darkened ellipses in A and the darkened boxes in B were
used in the weighted mean 206Pb / 238U zircon age calculation. U-Pb isotopic
data were not corrected for common lead (see text).

207Pb/206Pb, 207Pb / 235U, and 206Pb / 238U ages are reported for all twenty-five of the


WG05-02 zircon analyses in Table 4-3. Ages reported in Table 4-3 are given with 2C


errors. The 207Pb / 206Pb ages range from -1325-2804 Ma. However, all the WG05-02


zircon analyses gave discordant 207Pb / 235U and 206Pb / 238U ages except for zircon L8.


However, zircon L8 is only concordant because of the large errors associated with the









207Pb / 235U and 206Pb / 238U ages. The other twenty-four WG05-02 zircons yielded

discordant 207Pb / 235U and 206Pb / 238U ages ranging from -2476-285 Ma. Figure 4-3a

displays a traditional concordia diagram which includes all twenty-five of the WG05-02

zircon analyses. Twelve zircon analyses roughly defined a discordia line (see highlighted

ellipses). Figure 4-3b displays this discordia line which intersects concordia at 43 330

Ma and 1736 330 Ma 450 (2c errors). These intersection ages have large errors

because no zircon analyses fall near the intersections themselves.

As noted, all but one of the WG05-02 leucosome zircons (zircon L8) gave

discordant 207Pb / 235U and 206Pb / 238U ages. There are two possible explanations for the

largely discordant nature of the leucosome zircons: (1) One explanation is inheritance of

older zircon material from the leucosome zircon cores. Because the WG05-02 zircons

are small it is entirely possible that both core and rim regions of the zircons were

analyzed simultaneously (with beam diameter of 30-60 [tm). As a result, the 207Pb / 235U

and 206Pb / 238U zircon ages may represent a mixture of two zircon components (i.e., older

cores and younger rims). If different proportions of the two components were analyzed

from several zircons then a discordia array or mixing line could form when the isotopic

data are plotted on a concordia diagram (e.g., Figure 3a). (2) Another explanation for the

discordance observed in the WG05-02 leucosome zircons is lead loss from older,

probably Proterozoic or Archean zircons. For example, if very old zircons underwent

differential lead loss during a younger isotopic disturbance (e.g., subsequent reheating

during metamorphism and/or intrusion) these zircons would fall along a discordia line on

a concordia diagram. In this case, the two concordia-discordia intersects correspond to









the true age (upper intersect) of the old zircons and the age of younger isotopic

disturbance (lower intersect, Faure, 1986; Williams, 1998).

Both of these explanations require the presence of Proterozoic aged or older zircons

in the restite from which the WG05-02 leucosome was partially melted (i.e., the meta-

Greyson Fm). Ross and Villeneuve (2003) report 207Pb / 206Pb ages from detrital zircons

from unmetamorphosed Lower Belt Greyson-equivalent strata east (in the Helena Salient)

and west of the current study area that range from -1899-1670 Ma; these ages are

consistent with the 207Pb / 206Pb ages for the twelve WG05-02 zircon analyses that define

the discordia shown in figures 4-3a and b (-1525-1776 Ma, Fig. 4-4). Therefore, it is

possible the WG05-02 leucosome incorporated Mesoproterozoic or Paleoproterozoic

detrital zircons from the meta-Greyson Formation during high temperature

metamorphism and anatexis in the LISZ. Catholuminescence (CL) imaging of the

WG05-02 leucosome zircons is needed to determine which of the two above explanations

is correct. CL imaging can be used to determine if the leucosome zircons consist of

inherited cores with thin magmatic rims (consistent with explanation 1) or lack younger

magmatic rims (consistent with explanation 2). The latter implies that no new magmatic

rims grew on the older zircon cores during anatexis and formation of the granitic

leucosome.








74











0.0 . .----------------. Iy ---
a-ponmtorlip4ar. 2 2600 0

2200
04 O
180



no 0
""Pb 10-3 0 )

0.2
100, 0




00
0 2 4 8 8 10 12 14
207pbJ28U

0.4
diat0or1f eaMB es we 2

1800


140

'Pb 0.2
2mU 10


0.1 s6 Intercepts at 431 330 Ma
and 1736 450

200 B
0.0
0 2 4 6



Figure 4-3. Conventional U-Pb concordia plots for zircons for leucosome sample WG05-
02 collected from the meta-Greyson migmatitic paragneiss in the central
LISZ. A) Most leucosome zircons gave discordance ages with a 204Pb
common lead correction. B) Twelve leucosome zircon analyses fall along a
discordia that intersects concordia at 1736 450 Ma and 43 330. The
discordia is most likely the result of variable lead loss in Paleoproterozoic
zircons inherited from the meta-Greyson Formation protolith. Lead loss likely
occurred during anatexis and formation of the leucosome itself.

















0 X0


4-
Z 3

2-

0 I

1400 1800 2200 2600 3000 3400
207Pbj2Pb Age (Ma)
Figure 4-4. 207Pb/206Pb age density plot for the WG05-02 leucosome zircons. 207Pb/206Pb
ages range from -1525-3004 Ma. The leucosome zircons fall into two main
age groups: (1) a Mesoproterozoic to Paleoproterozoic group (-1525-1776
Ma) and an (2) Archean group (-2105-3004 Ma). The zircons that form the
discordia line shown in figures 3-4a and b are part of the Mesoproterozoic to
Paleoproterozoic group.














CHAPTER 5
THERMOBAROMETRY

Purpose and Strategy

As described above, the Lake of the Isle shear zone (LISZ) is a sinuous, middle to

upper-amphibolite facies ductile shear zone that stretches across the exhumed lower plate

of the ACC within the current study area (see Appendix F and description of the LISZ

above). Within the LISZ, Mesoproterozoic Belt Supergroup and middle Cambrian-

equivalent metasedimentary strata have undergone major ductile attenuation (thinning)

during upper amphibolite facies metamorphism. As a result, the deformed

metasedimentary strata of the LISZ exhibit a strongly transposed metamorphic foliation

and common mesoscopic foliation-parallel boudins and near isoclinal folds.

In the central LISZ, in the vicinity of the Lake of the Isle, migmatitic pelitic

paragneiss (correlated with metamorphosed Greyson Fm, Lower Belt) bears an

uppermost-amphibolite facies phase assemblage and shows evidence for anatexis during

the ductile deformation. In this part of the LISZ granitic leucosome is commonly found

in boudin necks and as isolated thin and elongate pods within the transposed foliation

indicating upper-amphibolite facies metamorphism and anatexis accompanied ductile

attenuation of Belt Supergroup and middle Cambrian-equivalent metasedimentary strata

in the central LISZ.

In the eastern part of the study area the LISZ is overprinted by the Eocene

Anaconda mylonite and several listric brittle normal faults related to the now largely

eroded brittle detachment system (Appendix F). Here, lower to middle greenschist facies









fabrics of the Anaconda granitoid mylonites contrast sharply with the plastically

deformed upper-amphibolite facies metasedimentary strata of the LISZ which lie

structurally beneath the mylonites. This structural relationship indicates that the LISZ

predated development of the Anaconda mylonite and brittle detachment system which

later facilitated the exhumation of the ACC lower plate. U-Pb geochronological and

40Ar/39Ar thermochronological data obtained from the LISZ in this study give a Late

Cretaceous age for the shear zone (ca. -75-74 Ma, see Chapter 4 and 6). These age

constraints indicate that upper-amphibolite facies metamorphism in the LISZ predated

tectonic exhumation of the ACC lower plate by probably no less than -23 Ma. The

structural relationship between the Anaconda greenschist facies mylonite zone and the

LISZ in the eastern study area also shows that the LISZ was exhumed from beneath the

brittle detachment fault system and ACC upper plate along with the greenschist mylonites

during the Eocene.

In this study, thermobarometric data were obtained from pelitic strata in the LISZ

to constrain the peak pressure-temperature associated with upper-amphibolite

metamorphism, local anatexis, and ductile attenuation of metasedimentary strata in the

shear zone. Because of the structural relationship between the LISZ and the greenschist

mylonites pressure constraints from the LISZ can be used to constrain the maximum

amount of exhumation facilitated by extension in the ACC.

ME-231 (Migmatitic Meta-Greyson Formation Paragneiss)

Sample ME-231 was selected for the thermobarometry needed in this study. ME-

231 was collected from cordierite-bearing migmatitic pelitic paragneiss (meta-Greyson)

exposed in the central part of the LISZ, just southwest of the Lake of the Isle (Appendix

F). Upper-amphibolite facies metamorphism was most intense here and the meta-









Greyson strata bear a phase assemblage of (in increasing relative abundance) cordierite +

garnet + K-feldspar + albite plagioclase + biotite + quartz + sillimanite fibrolite. This

metamorphic assemblage is indicative of uppermost amphibolite facies in the second

sillimanite-K-feldspar zone (Spear, 1993). A detailed description of ME-231 is included

in Appendix C.

Relevant Previous Pressure-temperature Constraints

Prior to this study, quantitative thermobarometric data were not available from the

metasedimentary strata deformed in the LISZ or any other part of the Anaconda-Pintlar

and Flint Creek Ranges. A few previous studies do document amphibolite facies regional

metamorphism of metasedimentary rocks in the ACC lower plate. The observations and

pressure-temperature estimates made in these studies are important to later discussion of

regional metamorphism in the ACC lower plate.

Desmarais (1983) described amphibolite facies pelitic and calc-silicate schists and

paragneisses in ACC lower plate exposed in the southern Anaconda-Pintlar Range,

southwest of the current study area. He described pelitic strata with an upper-amphibolite

facies assemblage of quartz + biotite + plagioclase + K-feldspar + sillimanite

muscovite. Desmarais correlated these strata with the Mount Shields Formation (middle

Missoula Group equivalent, upper Belt). He also observed migmatitic paragneisses in

some areas suggesting that local anatexis accompanied upper-amphibolite facies

metamorphism. Based on these observations, Desmarais estimated peak metamorphic

conditions in the southern Anaconda-Pintlar range reached pressures of -3.0-7.0 kbar and

temperatures of 600-7000C.

Flood (1974) mapped and described the Fishtrap Creek nappe, a large west-verging

recumbent nappe fold structure exposed in the southern Anaconda-Pintlar Range. The









Fishtrap Creek nappe is comprised of metasedimentary strata he correlated with (from

structurally lowest to highest) the Prichard Formation (Greyson Fm. equivalent), Ravalli

Group, and the Wallace Formation (Helena Fm. equivalent). Flood described these meta-

Belt units as biotite-muscovite schist, quartzite, and calc-silicate schist, respectively.

Flood (1974) also documented and measured well-developed metamorphic foliation

throughout the Fishtrap Creek nappe and common mesoscopic-scale folds. Based on the

metamorphic assemblages of these Belt-correlated metasedimentary strata Flood

estimated that peak metamorphic conditions in the southernmost Anaconda-Pintlar Range

reached pressures of -2-4 kbar and temperatures of -550-6500C.

Stuart (1966) mapped and described metasedimentary strata adjacent to the Late

Cretaceous Royal Stock granodiorite in the northeastern Flint Creek Range, north of the

current study area. He described a somewhat narrow contact aureole surrounding the

Royal Stock superimposed on more widespread and regional metamorphic fabrics.

Within the Royal Stock contact aureole he notes randomly oriented andalusite

porphyroblasts overprinting the regional metamorphic fabric. Outside the contact

aureole, Stuart documented highly deformed metasedimentary strata comprised of

muscovite + biotite + quartz + cordierite rather than andalusite. In thin section, he noted

large sigmoidal shaped cordierite porphyroblasts with their long axes oriented parallel to

the pervasive regional metamorphic foliation. These observations indicate an earlier

regional metamorphic event at middle to upper-amphibolite facies followed by a lower

pressure event associated with the intrusion of the Royal Stock. Notably, Stuart also

documented a general increase in the regional metamorphic grade west to east across the

northern Flint Creek Range, ranging from greenschist to upper-amphibolite facies.









Thermobarometry Results

Pressure-temperature estimates were obtained from polished thin sections prepared

from sample ME-231 by electron microprobe analyses and subsequent thermodynamic

calculations made using the computer programs AX (activity-composition) by Holland

and Powell (2000) and THERMOCALC v. 3.21 by Powell et al. (1998). For comparison,

a pressure-temperature estimate for ME-231 was also made using the computer program

Geothermobarometry (GTB) v. 2.1 by Spear and Kohn (1999). Sample preparation,

microprobe instrumentation and analytical procedures are summarized in Appendix A.

Thermodynamic calculations made using AX, THERMOCALC, and GTB are discussed

below.

Electron Microprobe Analyses

Major elemental compositions were obtained from individual mineral phases in

sample ME-231 using a JOEL Superprobe electron microprobe at Florida International

University. The results from the analyses of garnet, biotite, albite plagioclase and

cordierite are summarized in Table 5-1. Each mineral analysis is reported in elemental

oxide weight percent and element oxide totals from each analysis are shown in the far

right column of the table. Garnet and cordierite analyses labeled with the suffixes "c"

and "r", indicate mineral core and rim analyses, respectively. Garnet and cordierite

analyses without these suffixes were taken from intermediate spots on the mineral,

between the rim and core region. All biotite and plagioclase analyses from sample ME-

231 are rim analyses.

Garnet

A total of thirty-five spot analyzes were taken from large and small garnet

porphoroblasts. Most of these analyses gave elemental oxides totals of 100 1 weight














Table 5-1. Results from electron microprobe mineral analyses


Analyses Si02 TiO2 Cr203 Na20 K20 A1203 MnO FeO MgO CaO Total

Garnet
PMT-2-1-gtlc 36.94 0.23 0.17 0.32 0.10 21.29 1.06 36.93 2.54 1.02 100.60
PMT-2-1-gt2 37.36 0.15 0.17 0.37 0.10 21.46 1.29 36.18 2.97 1.03 101.06
PMT-2-1-gt3 37.02 0.12 0.14 0.33 0.10 21.33 1.28 36.30 2.99 1.09 100.70
PMT-2-1-gt4 37.00 0.23 0.25 0.31 0.11 21.50 1.04 36.62 2.99 0.94 100.97
PMT-2-1-gt5 36.72 0.21 0.15 0.28 0.11 21.29 0.58 37.85 2.24 1.14 100.57
PMT-2-3-gtl 36.68 0.05 0.06 0.01 0.01 21.05 0.33 37.72 2.47 1.22 99.61
PMT-2-3-gt2 36.42 0.07 0.07 0.00 0.01 21.29 0.44 37.50 2.56 0.90 99.26
PMT-2-3-gt3 36.51 0.00 0.06 0.02 0.00 21.19 0.40 36.59 2.50 1.12 98.40
PMT-2-3-gt4 36.41 0.00 0.05 0.00 0.01 21.36 0.36 36.86 2.55 0.95 98.54
PMT-2-6-gtl 36.37 0.00 0.04 0.00 0.00 21.08 0.29 38.31 2.13 0.82 99.03
PMT-2-6-gt2 36.12 0.00 0.02 0.00 0.00 21.13 0.49 34.29 2.25 1.19 95.50
PMT-2-6-gt3 36.79 0.00 0.00 0.00 0.01 21.07 0.49 37.68 2.18 1.00 99.21 -
PMT-3-1-gtl 36.81 0.13 0.15 0.30 0.13 21.57 0.51 38.35 2.91 1.10 101.94
PMT-3-1-gt2 36.73 0.13 0.16 0.31 0.12 21.39 1.13 36.78 3.11 0.95 100.81
PMT-3-1-gt3 36.60 0.18 0.14 0.29 0.13 21.24 1.40 36.75 2.89 1.02 100.64
PMT-3-1-gt4 36.17 0.13 0.16 0.29 0.12 21.14 0.95 36.64 2.94 1.04 99.57
PMT-3-1-gt5 36.58 0.17 0.12 0.29 0.11 21.38 0.49 37.96 2.93 1.26 101.28
PMT-3-1-gt6 36.70 0.21 0.33 0.30 0.11 21.13 0.76 37.25 2.93 1.45 101.16
PMT-3-1-gt7 36.18 0.12 0.34 0.27 0.12 21.17 0.81 36.84 2.54 1.30 99.70
PMT-3-1-gt8 36.37 0.08 0.16 0.29 0.13 20.93 0.63 37.50 2.97 1.25 100.30
PMT-3-1-gt9 36.54 0.19 0.26 0.31 0.12 21.58 0.99 36.94 2.96 1.29 101.17
PMT-3-1-gtl0 36.38 0.21 0.21 0.30 0.12 21.13 0.99 37.09 3.01 1.23 100.67














Table 5-1 Continued


Analyses Si02 TiO2 Cr203 Na20 K20 A1203 MnO FeO MgO CaO Total


PMT-3-2-gt1-1
PMT-3-2-gt1-2
PMT-3-2-gt1-3
PMT-3-2-gt2-1
PMT-3-2-gt2-2
PMT-3-2-gt2-3
PMT-3-3-gtl
PMT-3-3-gt2
PMT-3-3-gt3
PMT-3-3-gt4
PMT-3-3-gt5
PMT-3-3-gt6
PMT-3-3-gt7

Biotite
PMT-2-3-btl
PMT-2-3-bt2
PMT-2-3-bt3
PMT-2-6-btl
PMT-2-6-bt2
PMT-2-6-bt3
PMT-2-6-bt4
PMT-2-6-bt5


37.87
37.51
37.52
37.82
37.79
37.84
37.23
36.47
36.60
36.38
36.84
36.91
36.95


33.48
32.09
32.92
33.71
33.21
33.01
33.01
32.61


0.00
0.01
0.02
0.00
0.06
0.00
0.19
0.09
0.18
0.16
0.12
0.12
0.15


2.08
3.48
3.34
4.25
3.74
3.79
3.36
2.55


0.02
0.01
0.05
0.00
0.00
0.02
0.18
0.15
0.13
0.19
0.18
0.14
0.21


0.12
0.11
0.14
0.18
0.25
0.10
0.11
0.17


0.00
0.00
0.02
0.03
0.00
0.02
0.27
0.31
0.32
0.30
0.29
0.31
0.32


0.16
0.21
0.22
0.27
0.13
0.21
0.23
0.23


0.00 21.41
0.02 21.34
0.00 21.21


0.00
0.01
0.00
0.10
0.11
0.10
0.11
0.10
0.11
0.10


8.66
8.54
8.51
8.71
8.52
8.74
8.46
8.07


20.93
20.88
21.07
21.19
21.40
21.46
21.52
21.33
21.16
21.32


20.61
18.94
19.86
19.25
19.32
19.50
18.96
19.42


0.37 36.61
0.31 37.33
0.32 38.36
0.38 37.03
0.33 36.71
0.36 36.83
0.49 37.64
0.50 36.99
0.54 37.41
0.54 37.44
0.46 38.06
0.53 37.57
0.54 37.30


0.00
0.03
0.08
0.00
0.01
0.00
0.00
0.07


24.94
23.47
24.63
22.79
23.39
22.48
22.46
23.02


2.62
2.54
2.11
2.50
2.37
2.31
2.90
2.90
2.87
2.90
2.97
2.89
2.90


4.52
3.77
4.41
5.16
5.56
5.19
4.99
5.23


0.97
1.12
1.00
0.93
1.23
1.00
1.19
0.97
1.16
1.15
1.05
0.96
1.24


0.00
0.11
0.01
0.00
0.00
0.00
0.01
0.01


99.86
100.18
100.60
99.62
99.38
99.45
101.38
99.89
100.76
100.68
101.39
100.69
101.01


94.56
90.75
94.11
94.31
94.12
93.02
91.59
91.38














Table 5-1 Continued


Analyses Si02 TiO2 Cr203 Na20 K20 A1203 MnO FeO MgO CaO Total

PMT-2-6-bt6 32.58 3.92 0.06 0.25 8.47 18.96 0.05 21.25 5.06 0.01 90.61
PMT-2-6-bt7 32.74 4.46 0.10 0.23 7.93 19.29 0.00 22.46 4.82 0.00 92.04
PMT-3-1-btl 33.50 3.40 0.23 0.43 8.38 19.37 0.12 23.84 5.11 0.15 94.53
PMT-3-1-bt2 32.89 3.80 0.30 0.48 8.43 19.07 0.14 24.43 5.26 0.15 94.95
PMT-3-1-bt3 32.93 3.55 0.20 0.42 8.13 19.28 0.13 23.43 4.98 0.17 93.22
PMT-3-1-bt4 33.87 2.82 0.25 0.43 8.23 19.52 0.16 24.11 5.39 0.16 94.95
PMT-3-1-bt5 33.64 2.62 0.26 0.42 8.05 20.06 0.19 24.15 5.23 0.15 94.77
PMT-3-1-bt6 34.13 3.27 0.24 0.41 8.30 20.03 0.13 23.64 4.81 0.14 95.09
PMT-3-1-bt7 33.37 3.31 0.22 0.47 8.23 19.85 0.21 23.90 5.07 0.16 94.79
PMT-3-1-bt8 33.63 3.58 0.28 0.48 8.10 19.26 0.20 23.23 5.54 0.17 94.45
PMT-3-1-bt9 34.08 3.42 0.28 0.42 8.19 19.50 0.16 23.62 5.85 0.14 95.66
PMT-3-1-btl0 33.30 3.67 0.28 0.40 7.99 19.06 0.14 24.17 5.96 0.15 95.10
00
PMT-3-2-btl 33.92 2.02 0.06 0.21 8.70 20.85 0.07 24.19 4.58 0.01 94.60
PMT-3-2-bt2 34.05 2.85 0.15 0.20 8.60 20.05 0.00 24.03 4.55 0.00 94.47
PMT-3-2-bt3 33.78 2.58 0.05 0.18 8.54 19.83 0.07 24.48 4.84 0.01 94.35
PMT-3-2-bt4 33.67 2.84 0.20 0.16 8.93 19.46 0.00 23.56 4.97 0.00 93.81
PMT-3-2-bt5 34.08 2.52 0.13 0.16 8.61 19.72 0.00 23.62 5.07 0.00 93.91
PMT-3-2-bt6 33.58 3.14 0.10 0.19 8.81 19.98 0.02 22.82 5.06 0.01 93.71
PMT-3-3-btl 33.54 3.22 0.29 0.44 7.97 19.49 0.15 24.40 5.50 0.14 95.14
PMT-3-3-bt2 33.55 3.50 0.25 0.41 8.26 19.32 0.12 23.59 5.29 0.13 94.42
PMT-3-3-bt3 33.42 3.55 0.34 0.50 7.96 19.00 0.11 24.26 5.30 0.16 94.60
PMT-3-3-bt4 34.43 3.03 0.24 0.52 8.18 19.97 0.13 24.65 5.27 0.15 96.58
PMT-3-3-bt5 31.31 2.54 0.19 0.37 6.60 19.80 0.13 25.02 5.59 0.17 91.73














Table 5-1 Continued


Analyses Si02 TiO2 Cr203 Na20 K20 A1203 MnO FeO MgO CaO Total

Plagioclase
PMT-3-3-bt6 33.74 3.06 0.24 0.39 7.79 20.38 0.11 24.59 5.05 0.15 95.50
PMT-3-3-bt7 34.46 3.47 0.31 0.38 8.06 19.72 0.11 23.09 5.72 0.14 95.46
PMT-3-3-bt8 34.11 2.93 0.36 0.47 7.87 20.17 0.15 24.12 5.95 0.15 96.28
PMT-2-6-pl1 59.96 0.00 0.01 6.28 0.06 24.67 0.00 0.05 0.00 6.07 97.10
PMT-2-6-pl2 60.94 0.02 0.00 5.19 0.08 24.61 0.00 0.03 0.00 5.91 96.77
PMT-2-6-pl3 57.90 0.06 0.00 7.34 0.06 24.80 0.02 0.01 0.00 7.04 97.22
PMT-2-6-pl4 60.05 0.00 0.00 7.04 0.08 24.41 0.00 0.00 0.00 3.85 95.44
PMT-2-6-pl5 59.46 0.03 0.00 6.81 0.06 24.74 0.03 0.00 0.00 6.25 97.38
PMT-3-1-pl1 58.74 0.14 0.18 7.18 0.13 24.76 0.10 0.28 0.26 7.50 99.26
PMT-3-1-p12 58.57 0.08 0.07 7.18 0.17 24.65 0.11 0.42 0.27 7.58 99.10
PMT-3-1-p13 59.38 0.19 0.21 7.39 0.14 24.17 0.09 0.32 0.26 7.05 99.18
OO
PMT-3-1-p14 58.16 0.10 0.18 7.00 0.20 24.70 0.07 0.49 0.28 7.86 99.04 -
PMT-3-1-p15 59.02 0.05 0.13 7.12 0.13 24.54 0.10 0.34 0.28 7.38 99.07
PMT-3-1-p16 59.83 0.11 0.11 7.43 0.15 24.23 0.13 0.31 0.29 7.02 99.61
PMT-3-1-pl7 59.36 0.09 0.09 7.17 0.15 24.79 0.08 0.25 0.27 7.28 99.51
PMT-3-1-p18 60.35 0.11 0.14 7.63 0.13 23.99 0.09 0.31 0.28 6.53 99.55
PMT-3-2-pl1 61.02 0.00 0.05 7.38 0.06 23.97 0.01 0.23 0.00 6.41 99.12
PMT-3-2-pl2 60.55 0.00 0.00 7.52 0.05 24.11 0.00 0.13 0.00 6.47 98.84
PMT-3-2-pl3 60.94 0.00 0.08 7.70 0.06 23.95 0.00 0.22 0.00 6.40 99.35
PMT-3-2-pl4 61.86 0.00 0.01 5.39 0.05 24.34 0.02 0.18 0.00 6.49 98.35
PMT-3-3-pl1 59.52 0.00 0.10 6.37 0.04 25.47 0.05 0.11 0.00 8.10 99.75
PMT-3-3-pl2 58.89 0.00 0.00 6.12 0.03 25.18 0.03 0.09 0.00 8.30 98.64














Table 5-1 Continued


Analyses Si02 TiO2 Cr203 Na20 K20 A1203 MnO FeO MgO CaO Total


PMT-3-3-pl2
PMT-3-3-pl4
PMT-3-3-pl5
PMT-3-3-pl6
PMT-3-3-pl6
PMT-3-3-pl7
PMT-3-3-pl8
PMT-3-3-pl9


Cordierite
PMT-2-3-crdl
PMT-2-3-crd2
PMT-2-3-crd3
PMT-2-8-crdl
PMT-2-8-crd2
PMT-2-8-crd3
PMT-3-4-crdl
PMT-3-4-crd2
PMT-3-4-crd3


60.77
58.82
60.29
59.55
59.42
60.23
58.87
58.69


47.78
48.54
45.90
47.09
46.90
47.02
49.37
48.78
48.98


0.03
0.00
0.00
0.00
0.00
0.00
0.00
0.04


0.04
0.04
0.00
0.09
0.10
0.09
0.13
0.09
0.11


0.05
0.01
0.00
0.05
0.00
0.00
0.00
0.00


0.00
0.00
0.00
0.21
0.10
0.12
0.15
0.09
0.15


7.66
6.80
7.35
6.87
6.75
7.25
6.86
5.76


0.17
0.14
0.16
0.40
0.38
0.35
0.37
0.36
0.34


0.05
0.03
0.04
0.02
0.03
0.05
0.05
0.05


0.00
0.01
0.01
0.09
0.10
0.09
0.08
0.08
0.09


24.15
25.13
24.21
24.78
24.97
23.69
24.35
24.15


31.02
31.35
32.55
31.95
32.28
31.73
31.75
31.32
31.21


0.00
0.04
0.01
0.00
0.03
0.02
0.01
0.03


0.07
0.07
0.00
0.09
0.15
0.09
0.17
0.13
0.16


0.31
0.29
0.27
0.11
0.12
0.05
0.29
0.44


12.16
11.13
11.45
11.48
10.76
11.66
11.05
11.02
11.22


0.00
0.02
0.00
0.00
0.00
0.00
0.07
0.00


5.23
6.29
6.30
6.53
6.84
6.46
6.64
6.91
6.70


6.52
8.12
7.10
7.78
7.86
6.88
7.07
6.81


0.03
0.03
0.02
0.15
0.13
0.13
0.13
0.15
0.15


99.52
99.26
99.26
99.15
99.18
98.17
97.58
95.97


96.50
97.60
96.39
98.08
97.74
97.74
99.85
98.92
99.10


Note: gt = garnet, bt = biotite, pl = plagioclase, crd = cordierite. The letter c indicates a core analyses. All analyses are reported in
weight percent elemental oxides.