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Precambrian Crustal Evolution in the Great Falls Tectonic Zone

Permanent Link: http://ufdc.ufl.edu/UFE0045248/00001

Material Information

Title: Precambrian Crustal Evolution in the Great Falls Tectonic Zone
Physical Description: 1 online resource (212 p.)
Language: english
Creator: Gifford, Jennifer N
Publisher: University of Florida
Place of Publication: Gainesville, Fla.
Publication Date: 2013

Subjects

Subjects / Keywords: geochemistry -- geochronology
Geological Sciences -- Dissertations, Academic -- UF
Genre: Geology thesis, Ph.D.
bibliography   ( marcgt )
theses   ( marcgt )
government publication (state, provincial, terriorial, dependent)   ( marcgt )
born-digital   ( sobekcm )
Electronic Thesis or Dissertation

Notes

Abstract: The Great Falls Tectonic Zone (GFTZ) is a zone of northeast trending geological structures in central Montana that parallel structures in the underlying basement.  U-Pb zircon and Nd isotopic data from the Little Belt Mountains (LBM) suggest that the GFTZ formed at ~1.86 to 1.80 Ga due to ocean subduction followed by collision between the Archean Wyoming Province (WP) and Medicine Hat Block (MHB).  This study characterizes the GFTZ basement by geochronological and geochemical analysis of crustal xenoliths collected from Montana Alkali Province volcanics and exposed basement rock in the Little Rocky Mountains (LRM).   Xenoliths collected from the Grassrange and Missouri Breaks diatremes and volcanics in the Bearpaw and Highwood Mountains have igneous crystallization ages from ~1.7 Ga to 1.9 Ga and 2.4 Ga to 2.7 Ga, and metamorphic ages from ~1.65 Ga to 1.84 Ga.  Zircon Lu-Hf and whole-rock Sm-Nd data indicate that the xenoliths originated from reworked older continental crust mixed with mantle-derived components in all cases.  Trace element patterns show fluid mobile element enrichments and fluid immobile element depletions suggestive of a subduction origin.  Igneous ages in the LRM range older, from ~2.4 Ga to 3.2 Ga.  Geochemical evidence suggests that the LRM meta-igneous units also formed in a subduction setting.  Detrital zircon ages span the early Paleoproterozoic to Mesoarchean, with abundant 2.8 Ga ages. Zircon U-Pb igneous crystallization age data from xenoliths and the LRM are consistent with U-Pb zircon igneous crystallization ages from the MHB, suggesting that this segment of the GFTZ shares an affinity with concealed MHB crust.  Published detrital zircon ages from the northern Wyoming Province reveal more abundant >3.0 Ga ages than the MHB or GFTZ samples.  These geochronologic and geochemical data from the xenoliths and LRM samples allow for a refined model for crustal evolution in the GFTZ.  Subduction under the Neoarchean to Paleoproterozoic crust of the MHB formed an igneous arc followed by metamorphism during the MHB-WP collision.  Later Paleoproterozoic tectonothermal activity represents post-orogenic collapse after the terminal collision.  Tectonic activity in the Cenozoic led to basement uplift and the formation of xenolith bearing volcanic units sampled for this study.
General Note: In the series University of Florida Digital Collections.
General Note: Includes vita.
Bibliography: Includes bibliographical references.
Source of Description: Description based on online resource; title from PDF title page.
Source of Description: This bibliographic record is available under the Creative Commons CC0 public domain dedication. The University of Florida Libraries, as creator of this bibliographic record, has waived all rights to it worldwide under copyright law, including all related and neighboring rights, to the extent allowed by law.
Statement of Responsibility: by Jennifer N Gifford.
Thesis: Thesis (Ph.D.)--University of Florida, 2013.
Local: Adviser: Foster, David A.
Local: Co-adviser: Mueller, Paul A.

Record Information

Source Institution: UFRGP
Rights Management: Applicable rights reserved.
Classification: lcc - LD1780 2013
System ID: UFE0045248:00001

Permanent Link: http://ufdc.ufl.edu/UFE0045248/00001

Material Information

Title: Precambrian Crustal Evolution in the Great Falls Tectonic Zone
Physical Description: 1 online resource (212 p.)
Language: english
Creator: Gifford, Jennifer N
Publisher: University of Florida
Place of Publication: Gainesville, Fla.
Publication Date: 2013

Subjects

Subjects / Keywords: geochemistry -- geochronology
Geological Sciences -- Dissertations, Academic -- UF
Genre: Geology thesis, Ph.D.
bibliography   ( marcgt )
theses   ( marcgt )
government publication (state, provincial, terriorial, dependent)   ( marcgt )
born-digital   ( sobekcm )
Electronic Thesis or Dissertation

Notes

Abstract: The Great Falls Tectonic Zone (GFTZ) is a zone of northeast trending geological structures in central Montana that parallel structures in the underlying basement.  U-Pb zircon and Nd isotopic data from the Little Belt Mountains (LBM) suggest that the GFTZ formed at ~1.86 to 1.80 Ga due to ocean subduction followed by collision between the Archean Wyoming Province (WP) and Medicine Hat Block (MHB).  This study characterizes the GFTZ basement by geochronological and geochemical analysis of crustal xenoliths collected from Montana Alkali Province volcanics and exposed basement rock in the Little Rocky Mountains (LRM).   Xenoliths collected from the Grassrange and Missouri Breaks diatremes and volcanics in the Bearpaw and Highwood Mountains have igneous crystallization ages from ~1.7 Ga to 1.9 Ga and 2.4 Ga to 2.7 Ga, and metamorphic ages from ~1.65 Ga to 1.84 Ga.  Zircon Lu-Hf and whole-rock Sm-Nd data indicate that the xenoliths originated from reworked older continental crust mixed with mantle-derived components in all cases.  Trace element patterns show fluid mobile element enrichments and fluid immobile element depletions suggestive of a subduction origin.  Igneous ages in the LRM range older, from ~2.4 Ga to 3.2 Ga.  Geochemical evidence suggests that the LRM meta-igneous units also formed in a subduction setting.  Detrital zircon ages span the early Paleoproterozoic to Mesoarchean, with abundant 2.8 Ga ages. Zircon U-Pb igneous crystallization age data from xenoliths and the LRM are consistent with U-Pb zircon igneous crystallization ages from the MHB, suggesting that this segment of the GFTZ shares an affinity with concealed MHB crust.  Published detrital zircon ages from the northern Wyoming Province reveal more abundant >3.0 Ga ages than the MHB or GFTZ samples.  These geochronologic and geochemical data from the xenoliths and LRM samples allow for a refined model for crustal evolution in the GFTZ.  Subduction under the Neoarchean to Paleoproterozoic crust of the MHB formed an igneous arc followed by metamorphism during the MHB-WP collision.  Later Paleoproterozoic tectonothermal activity represents post-orogenic collapse after the terminal collision.  Tectonic activity in the Cenozoic led to basement uplift and the formation of xenolith bearing volcanic units sampled for this study.
General Note: In the series University of Florida Digital Collections.
General Note: Includes vita.
Bibliography: Includes bibliographical references.
Source of Description: Description based on online resource; title from PDF title page.
Source of Description: This bibliographic record is available under the Creative Commons CC0 public domain dedication. The University of Florida Libraries, as creator of this bibliographic record, has waived all rights to it worldwide under copyright law, including all related and neighboring rights, to the extent allowed by law.
Statement of Responsibility: by Jennifer N Gifford.
Thesis: Thesis (Ph.D.)--University of Florida, 2013.
Local: Adviser: Foster, David A.
Local: Co-adviser: Mueller, Paul A.

Record Information

Source Institution: UFRGP
Rights Management: Applicable rights reserved.
Classification: lcc - LD1780 2013
System ID: UFE0045248:00001


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1 PRECAMBRIAN CRUSTAL EVOLUTION IN THE GREAT FALLS TECTONIC ZONE By JENNIFER N. GIFFORD A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGRE E OF DOCTOR OF PHILOSOPHY UNIVERSITY OF FLORIDA 2013

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2 2013 Jennifer N. Gifford

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3 To my family and friends, without your support I never would have made it this far.

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4 ACKNOWLEDGMENTS I wish to thank the National Science Foundation for funding my project through grants EAR0545715, EAR0538133 EAR0546751 and EAR0609952. I wish to thank my advisors Dr. Paul Mueller and Dr. David Foster for their guidance and patience. I also wish to thank my committee ; Dr. Karen Bjorndal, Dr. Kyle Min, Dr. Michael Perfit, and Dr. Ray Russo, for their guidance and helpful reviews of my dissertation material. I wish to thank my family for their un ending love and support. And finally, I wish to thank Dr. Shawn Malone for hi s support and assistance.

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5 TABLE OF CONTENTS page ACKNOWLEDGMENTS ................................ ................................ ................................ .. 4 LIST OF TABLES ................................ ................................ ................................ ............ 7 LIST OF FIGURES ................................ ................................ ................................ .......... 9 LIST OF OB JECTS ................................ ................................ ................................ ....... 15 LIST OF ABBREVIATIONS ................................ ................................ ........................... 16 ABSTRACT ................................ ................................ ................................ ................... 17 CHAPTER 1 INTRODUC TION ................................ ................................ ................................ .... 19 2 GRASSRANGE XENOLITHS ................................ ................................ ................. 25 Introduction ................................ ................................ ................................ ............. 25 Geologic Backg round ................................ ................................ .............................. 26 Results ................................ ................................ ................................ .................... 27 Whole Rock Geochemistry ................................ ................................ ............... 28 U Pb Geochronolog y In Zircon ................................ ................................ ......... 30 Hf Isotopes In Zircon ................................ ................................ ........................ 31 Sm Nd Whole Rock Isotopes ................................ ................................ ........... 33 Discussion ................................ ................................ ................................ .............. 34 Origins Of The Granitic Xenoliths ................................ ................................ ..... 34 Origins Of The Schistose and Quartzite Xenoliths ................................ ........... 36 Implications For Great Falls Tectonic Zone Evolution ................................ ...... 38 Conclusions ................................ ................................ ................................ ............ 41 3 MISSOURI BR EAKS XENOLITHS ................................ ................................ ......... 70 Introduction ................................ ................................ ................................ ............. 70 Geologic Background ................................ ................................ .............................. 72 Sample Descriptions ................................ ................................ ............................... 74 Big Slide Diatreme ................................ ................................ ............................ 74 Robinson Ranch Diatreme ................................ ................................ ............... 74 Little Sand Creek Diatreme ................................ ................................ .............. 75 Bearpaw Mountains At Lloyd Divide ................................ ................................ 76 Highwood Mountains ................................ ................................ ........................ 77 Results ................................ ................................ ................................ .................... 77 Whole Rock Geochemistry ................................ ................................ ............... 77 U Pb Geochronology Of Zircon ................................ ................................ ........ 78

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6 Hf Isotopes In Zircon ................................ ................................ ........................ 84 Sm Nd Whole Rock Isotopes ................................ ................................ ........... 88 Discussion ................................ ................................ ................................ .............. 90 Origins Of The Meta Igneous Xenoliths ................................ ........................... 90 Implications For Great Falls Tectonic Zone Evolution ................................ ...... 93 Conclusions ................................ ................................ ................................ ............ 97 4 LITTLE ROCKY MOUNTAINS ................................ ................................ .............. 130 Introduction ................................ ................................ ................................ ........... 130 Geologic Background ................................ ................................ ............................ 131 Stratigraphy ................................ ................................ ................................ .... 132 Previous Geochronology And Geochemistry ................................ .................. 134 Sample Descriptions ................................ ................................ ............................. 134 Amphibolite ................................ ................................ ................................ ..... 134 Quartzofeldspathic Schists ................................ ................................ ............. 135 Gneisses ................................ ................................ ................................ ........ 136 Other Samples ................................ ................................ ............................... 137 Results ................................ ................................ ................................ .................. 138 Whole Rock Geochemistry ................................ ................................ ............. 138 U Pb Geochronology Of Zircon ................................ ................................ ...... 140 Hf Isotopes In Zircon ................................ ................................ ...................... 144 Sm Nd Whole Rock Isotopes ................................ ................................ ......... 147 Discussion ................................ ................................ ................................ ............ 148 Geochemical Insight Into Sample Origins ................................ ....................... 148 Little Rocky Mountains as Exposed Medicine Hat Block Crust ....................... 149 Correlation of igneous ages ................................ ................................ ..... 149 Insights into meta sediment provenance ................................ ................. 150 Other regional considerations ................................ ................................ .. 153 Conclusions ................................ ................................ ................................ .......... 155 5 SUMMARY/CONCLUSIONS ................................ ................................ ................ 187 APPENDIX A METHODS ................................ ................................ ................................ ............ 195 U Pb And Hf Isotopic Analy sis Of Zircon ................................ .............................. 195 Whole Rock Geochemistry ................................ ................................ ................... 196 B SUPPLEMENTARY DATA TABLES ................................ ................................ ..... 198 LIST OF REFERENCES ................................ ................................ ............................. 199 BIOGRAPHICAL SKETCH ................................ ................................ .......................... 212

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7 LIST OF TABLES Table page 2 1 Major, trace and rare earth element data of granitoid xenoliths from the Grassrange. Major elements in wt. %, trace and rare earth elements in ppm ... 43 2 2 Major, trace and rare earth e lement data of schist and quartzite xenoliths from the Grassrange. Major elements in wt. %, trace and rare e arth elements in ppm ................................ ................................ ................................ 45 2 3 Granitoid xenoliths, Grassrange, zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf isotope data reported (Ma) ................................ ............................. 47 2 4 Schist and quartzite xenoliths, Grassrange, zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf isotope data reported (Ma) ................................ .......... 48 2 5 LA ICP M S Nd isotope data reported (Ma) ................................ ........................ 49 3 1 Major, trace and rare earth element data of igneous protolith xenoliths from the Big Slide and Little Sand Creek diatremes. Major elements in wt. %, trace and rare earth elements in ppm ................................ ................................ .......... 99 3 2 Major, trace and rare earth element data of xenoliths from Robinson Ranch Diatre me, the Bearpaw Mountains, and the Highwood Mountains. Major elements in wt. %, trace and rare earth ele ments in ppm ................................ 101 3 3 Igneous protolith xenoliths, Missouri Breaks, zircon LA ICP MS U Pb data, 207 Pb/ 206 Pb ages an d Hf isotope data reported (Ma) ................................ ........ 103 3 4 LA ICP M S Nd isotope data reported (Ma) ................................ ...................... 104 4 1 Major, trace a nd rare earth element data of orthogneiss samples from the Little Rocky Mountains. Major elements in wt. %, trace and rare earth elements in ppm ................................ ................................ ............................... 157 4 2 Major, trace and rare earth element da ta of amphibolite samples from the Little Rocky Mountains. Major elements in wt. %, trace and rare earth elements in ppm ................................ ................................ ............................... 159 4 3 Major, trace and rare earth element data of schist samples fro m the Little Rocky Mountains. Major elements in wt. %, trace and rare earth elements in ppm ................................ ................................ ................................ .................. 161 4 4 Major, trace and rare earth element data of paragneiss samples from the Little Rocky Mounta ins. Major elements in wt. %, trace and rare earth elements in ppm ................................ ................................ ............................... 163

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8 4 5 Orthogneiss and amphibolite samples, Little Rocky Mountains, zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf isot ope data reported (Ma) .......... 165 4 6 Schist and paragneiss samples, LRM, zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages an d Hf isotope data reported (Ma) ................................ ........ 166 4 7 Whole rock LA ICP M S Nd isotope data reported (Ma) ................................ .... 167

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9 LIST OF FIGURES Figure page 1 1 (A) Paleoproterozoic Laurentia (aft er Hoffman, 1988; Foster et al., 2006; Davidson, 2008) and a location map showing study area. (B) Generalized map of Precambrian basement provinces of southwestern Laurentia ................ 23 1 2 Generalized depiction of Cenozoic alkaline rock occurrences in the Montana alkali province (after Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well defined and are based on ................. 24 2 1 Generalized map of Precambrian basement provinces of southwestern Laurentia. Exposures of basement in Laramide style uplifts are show n in the dark grey shaded areas ................................ ................................ ...................... 50 2 2 Generalized depiction of Cenozoic alkaline rock occurrences in the Montana alkali province (after Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well defined ................................ ............. 51 2 3 Representative examples of photomicrographs of (A) granitic sample MX 08 in plain polarized light (PPL); (B) granitic sample MX 08 in cross polarized light (XPL); (C) schist sample HAL 1 in PPL; (D) schist sample HAL 1 .............. 52 2 4 Granitoid samples plotted on an alumina saturation index (Shand, 1943; modified by Frost et al., 2001). ~2.5 Ga granitoids (diamonds ), ~1.7 Ga granitoids (squares) ................................ ................................ ............................ 53 2 5 Plot of P 2 O 5 /TiO 2 vs. MgO/CaO (after Werner, 1987) proposed to discriminate magmatic vs. sedimentary protoliths for all meta igneous g ranitoids and schistose samples ................................ ................................ ....... 54 2 6 Compositional discriminant diagram showing the discriminant function 3 (DF3) vs. SiO 2 for all granitoid and schistose samples (after Shaw, 1972). Rocks with positive DF3 values are interpreted to be of igneous origin .............. 55 2 7 Primitive for (A) granitoid samples and (B) schistose samples, using data in Table 2 1 and 2 2. Relative enrichment of large ion lithophile element s ............................ 56 2 8 Chondrite normalized rare earth element diagram (McDonough and Sun, 1995) for granitoid samples using data in Table 2 1 ................................ ........... 57 2 9 Concordia diagrams showing U Pb data for granitoid samples MX 08 (A) and MX 18 (B). Each ellipse represents a single spot ana standard error ................................ ................................ ................................ ..... 58

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10 2 10 (A) Concordia diag rams plotting composite U Pb data for granitoid sample MX error. (B) Expanded view of lower end of concordia showing samples .............. 59 2 11 (A) Concordia diagrams plotting composite U Pb data for granitoid sample MX error. (B) Expanded view of lower end of concordia showing samples .............. 60 2 12 (A) Concordia diagrams plotting composite U Pb data for granitoid sample MX error. (B) Expanded view of lower end of concordia showing sa mples .............. 61 2 13 (A) Concordia diagrams plotting composite U Pb data for granitoid sample MX error. (B) Expanded view of l ower end of concordia showing samples .............. 62 2 14 Probability density plots comparing the cumulative <10% discordant zircon analyses of (A) the 6 granitoid samples, 155 grains; (B) The 5 schist sa mples, 59 grains; (C) The 3 quartzite samples, 95 grains ............................... 63 2 15 Sm Nd evolution diagram showing the range for the ~1.7 granitoids (gray stars), ~2.5 granitoids (white stars), Little Belt moun tains (white bar), and the generalized evolution of northern Wyoming province crust ................................ 64 2 16 Histogram showing Sm Nd depleted mantle model ages (Ga) of granitoid, schist, and quartzite whole ro cks. Depleted Mantle model ages calculated u sing the model of DePaolo, 1981 ................................ ................................ ...... 65 2 17 Hf evolution diagram showing the range for the ~1.7 granitoids (circles), ~2.5 granitoids (diamonds), schist ose (square), quartzite (triangle), and LBMs (black dashes) from ages 1.70 Ga to 1.89 Ga ................................ .................... 66 2 18 Published Archean to earliest Proterozoic U Pb ages of zircon from the northern Wyoming craton compared to published, ages for the MHB. OC Owl Creek Mountains, BM Bighorn Mountains ................................ ................. 67 2 19 Trace element discrimination diagrams (after Pearce et al., 1984) for granitoid samples: (A) H eavy rare earth element Y ppm vs. high field strength element Nb ppm, (B) Heavy rare earth element Yb ppm ................................ .... 68 2 20 Comparison of published U Pb ages from the southern Trans Hudson Orogen again st the northern Trans Hudson Orogen, Great Falls tectonic zone (this study), and Yavapai Central Plains ................................ .................... 69 3 1 Generalized map of Precambrian basement provinces of southwestern Laurentia (afte r Ross et al., 1991; Condie, 1992; Doughty et al., 1998; Vogl et al., 2004; Foster et al., 2006, 2012). Exposures of basement ..................... 105

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11 3 2 Generalized depiction of Cenozoic alkaline rock occurrenc es in the Montana alkali province (after Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well defined ................................ ........... 106 3 3 Photomicrographs of selected meta igneou s samples, with Plane Polarized light (PPL) images on the left, and Cross Polarized Light (XPL) images on the right. Scale bar is 1 mm. (A) Sample LD10 01, a garnet .......................... 107 3 4 Photomicrographs of selected meta igneous samples, with Plane Polarized light (PPL) images on the left, and Cross Polarized Light (XPL) images on the right. Scale bar is 1 mm. (A) BSD10 05 is a mafic meta granitoid ............ 108 3 5 Xenolith samples plotted on an alumina saturation index (Shand, 1943; modified by Frost et al., 2001) Meta igneous blue diamonds ....................... 109 3 6 (A) Primitive 1995) for meta igneous samples. (B) Chondrite normalized REE diagram (McDonough and Sun, 1995 ) for meta igneous samples ................................ 110 3 7 (A) Concordia diagram plotti ng composite U Pb data for meta granitoid sample BSD10 error. (B) Expanded view of lower end of concordia ................................ ........ 111 3 8 Concordia diag ram plotting composite U Pb data for mafic meta granitoid sample BSD10 05. Each ellipse represents a single error ................................ ................................ ................................ .................. 112 3 9 Concordia diagram showing U Pb upper interce pt regression for amphibolite samples LSC10 03 (A) and LSC10 10 (B) as well as mafic granulitic gneiss LSC10 13 (C). Each ellipse represents a single spot analysis ........................ 113 3 10 (A) Concordia diag Pb data for mafic gneiss sample LSC10 11. Each ellipse represents a single spot analysis and ................ 114 3 11 Concordia diagram showing U Pb upper intercept regression for granitoid samples RRD10 05 (A) and RRD10 09 (B). Each ellipse represents a single ................................ ................................ ........... 115 3 12 (A) Concordia diagram plotting composite U Pb data for garnet granulite sample RRD10 error. (B) Expanded view of upper end of concordia showing calculated ........ 116 3 13 Concordia diagram plotting composite U Pb discordia regression for orthogneiss sample LD10 07. Each ellipse represents a single spot analysis ................................ ................................ ................................ 117

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12 3 14 (A) Concordia diagram plotting composite U Pb data for orthogneiss sample LD10 Expanded view of lower end of concordia showing spread of samples ............ 118 3 15 (A) Concordia diagram plotting composite U Pb data for orthogneiss sample LD10 Expanded view of lower end of concordia showing spread of samples ............ 119 3 16 Concordia diagram showing U Pb upper intercept regression for quartz pegmatite sample LD10 11. Each ellipse represents a single spot analysis ................................ ................................ ................................ 120 3 17 (A) Concordia diagrams plotting composite U Pb data for orthogneiss sample HX Expanded view of upper end of conco rdia showing samples ........................... 121 3 18 (A) Concordia diagram plotting composite U Pb data for dioritic granulite sample RRD10 20. (B) Concordia diagram plotting composite U Pb data for meta granitoid sample LSC10 12. Each ellipse represents a single ................ 122 3 19 Hf evolution diagram showing the granitoid and meta igneous samples (blue diamonds) against the granitoid samples from the Grassrang e (Chapter 2) (orange squares) and data from the Little Belt Mountains ................................ 123 3 20 07. (B) Probability density plot showing Hf T (DM) values for sample LD10 07. Met. metamorphic zircons; Mag. magmatic zircons ................................ .... 124 3 21 Sm Nd evolution diagram showing the meta igneous samples (blue diamonds), Little Belt Mountains (green bar), and four different northern Wyoming province crustal outcrops (all citations are in the text) ...................... 125 3 22 Histogram showing Sm Nd depleted mantle model ages (Ga) of meta igneous whole rocks. T DM calculated u sing the model of DePaolo, 1981 ......... 12 6 3 23 Trace elem ent discrimination diagrams (after Pearce et al., 1984) for meta igneous samples (blue diamonds): (A) Heavy rare earth element Y (ppm) vs. high field strength element Nb (ppm), (B) Heavy rare earth element Yb .......... 127 3 24 Published Archean to earliest Paleoproterozoic U Pb ages of zircon from the northern Wyoming craton (red bars) compared to published ages for the MHB (blue bars) (all citations are in the text). OC Owl Creek Mountains ...... 128 3 25 Comparison of published U Pb ages from the southern Trans Hudson Orogen (all citations are in the text) against the northern Trans Hudson Orogen, Great Falls tectonic zone (Chapter 2 and this s tudy), and .................. 129

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13 4 1 Generalized map of Precambrian basement provinces of southwestern Laurentia (after Ross et al., 1991; Condie, 1992; Doughty et al., 1998; Vogl et al., 2004; Foster et al., 2 006, 2012). Exposures of basement ..................... 168 4 2 Location of the Little Rocky Mountains relative to a generalized depiction of Cenozoic alkaline rock occurrences in the Montana alkali province (afte r Hearn et al., 1989). The limits of the Great Falls tectonic zone ....................... 169 4 3 Map of the Little Rocky Mountains (after Hearn et al., 1989). Sample loca tions are shown as yellow stars ................................ ................................ 170 4 4 Photomicrographs of selected meta plutonic samples from the Little Rocky Mountains. Mineral abbreviations are after Whitney and Evans (2010). Plane light (PPL) mircrographs are on the left ................................ .................. 171 4 5 Photomicrographs of selected meta supracrustal samples from the Little Rocky Mountains. Mineral abbreviations are after Whitney and Evans (2010). PPL mircrographs are on the left, and X PL micrographs .................... 172 4 6 Schist, orthogneiss and paragneiss samples plotted on an alumina saturation index (Shand, 1943; modified by Frost et al., 2001). Schists (green triangles), orthogneisse s (blue diamonds), and paragneisses .......................... 173 4 7 Ternary CaO MgO Al2O3 variation diagram showing the expected fields of meta igneous and meta sedimentary rocks for all schists (circles) and paragne isses (squares) (after Leyreloup et al., 1977) ................................ ...... 174 4 8 Primitive for (A) orthogneiss and amphibolite samples and (B) detrital schi stose and paragneiss samples. Relative enrichment of large ion lithophile ...................... 175 4 9 Chondrite normalized diagram (McDonough and Sun, 1995) for orthogneiss and amphibolite samples using data from Tables 4 1 and 4 2. Orthogneiss samples show elevated concentrations of light versus ................................ ..... 176 4 10 Concordia diagrams showing U Pb data showing spread of zircons and weighted mean age for orthogn eiss samples MLR 01 (A) and MLR 06 (B). Each ellipse represents a single spot analysis ........................ 177 4 11 Concordia diagrams showing U Pb data for orthogneiss samples MLR 15 (A) and MLR 19 (B). Each ellipse represents a single error ................................ ................................ ................................ .................. 178 4 12 (A) Concordia diagrams plotting composite U Pb data for orthogneiss sample LRM Expanded view of upper end of concordia sho wing the oldest, most ............... 179

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14 4 13 Concordia diagram showing U Pb data for orthogneiss sample LRM 5. Each ellipse represents a single ................................ 180 4 14 Concordia diagram showing U Pb data for orthogneiss sample LRM 7. Each ellipse represents a single ................................ 181 4 15 (A) Conco rdia diagrams plotting composite U Pb data for orthogneiss sample MLR 09. Red dashed lines indicate trends within the analyses. (B) Expanded view of upper end of concordia showing the oldest ......................... 182 4 16 (A) Concordia diagrams plotting composite U Pb data for amphibolite sample MLR 03. Red dashed lines indicate trends within the analyses. (B) Expanded view of upper end of concordia showing the oldest ......................... 183 4 17 analyses of the schist and paragneiss samples. Orthogneiss ages are included for comparison ................................ ................................ ................... 184 4 18 (A) Lu Hf evolution diagram showing the range for the orthogneisss (blue diamond), amphibolites (orange circle), schistose (green triangle), and paragneisses (red square). (B) Sm Nd evolution diagram ............................... 185 4 19 Published Archean to earliest Proterozoic U Pb ages of zircon from the northern Wyoming craton (citations in the text) compared to published, ages for the MHB (citations in text). Wyoming yellow bars ................................ ..... 186 5 1 Schematic evolution of the Great Falls tectonic zone based on geologic constraints and chronology. (A) map view: ocean subduction between the Wyoming province (WP) and the Medicine Hat Block (MHB) ........................... 194

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15 LIST OF OBJECTS Object page B 1 External sup plemental data tables ................................ ................................ .... 198

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16 LIST OF ABBREVIATIONS BDL B elow detection limit BSE Bulk silicate earth BTM Beartooth Mountains CCB ULK Bulk continental crust CHUR Chondritic uniform reservoir DM Depleted mantle GFTZ Great Falls tectonic zone HFSE High field strength element HREE Heavy rare earth element LBM Littl e Belt Mountains LIL Large ion lithophile LOI Loss on ignition LRM Little Rocky Mountains MHB Medicine Hat Block MR Madison Range MSWD Mean standard weighted deviates N/A Not applicable PM Primitive mantle P PM Parts per million REE Rare earth elements THO Trans Hudson Orogeny TR Teton Range TRM Tobacco Root Mountains W T % Weight percent

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17 Abstract of Dissertation Presented to the Graduate School of the University of Florida in Partial Fulfillment of the Requirements for the Degree of Doctor of Philosophy PRECAMBRIAN CRUSTAL EVOLUTION IN THE GREAT FALLS TECTONIC ZONE By Jennifer N. Gifford May 2013 Chair: David A. Foster Cochair: Paul A. Mueller Major: Geological Sciences The Great Falls Tectonic Zone (GFTZ) is a zone of northeast trendi ng geological structures in central Montana that parallel structures in the underlying basement. U Pb zircon and Nd isotopic data from the Little Belt Mountains (LBM) suggest that the GFTZ form ed at ~1.86 to 1.80 Ga due to ocean subduction followed by col lision between the Archean Wyoming Province ( W P) and Medicine Hat Block ( MHB ) This study characteriz es the GFTZ basement by geochronolog i cal and geochemical analysis of crustal xenoliths collected from Montana Alkali Province volcanics and exposed baseme nt rock in the Little Rocky Mountains (LRM). Xenoliths collected from the Grassrange and Missouri Breaks diatremes and volcanics in the Bearpaw and Highwood Mountains have igneous crystallization ages from ~1.7 Ga to 1.9 Ga and 2.4 Ga to 2.7 Ga, and meta morphic ages from ~1.65 Ga to 1.84 Ga. Zircon Lu Hf and whole rock Sm Nd data indicate that the xenoliths originated from reworked older continental crust mixed with mantle derived components in all cases. Trace element patterns show fluid mobile element enrichments and fluid immobile element depletions suggestive of a subduction origin. Igneous ages in the

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18 LRM range older, from ~2.4 Ga to 3.2 Ga. Geochemical evidence suggests that the LRM meta igneous units also formed in a subduction setting. Detrita l zircon ages span the early Paleoproterozoic to Mesoarchean, with abundant 2.8 Ga ages. Zircon U Pb igneous crystallization age data from xenoliths and the LRM are consistent with U Pb zircon igneous crystallization ages from the MHB, suggesting that this segment of the GFTZ shares an affinity with concealed MHB crust. Published detrital zircon a ges from the northern Wyoming Province reveal more abundant >3.0 Ga ages than the MHB or GFTZ samples These g eochronologic and geochemical data from the xenolit hs and LRM samples allow for a refined model for crustal evolution in the GFTZ. S ubduction under the Neoarchean to Paleoproterozoic crust of the MHB formed an igneous arc followed by metamorphism during the MHB W P collision. Later Paleoproterozoic tecton othermal activity represents post orogenic collapse after the terminal collision. Tectonic activity in the Cenozoic led to basement uplift and the formation of xenolith bearing volcanic units sampled for this study.

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19 CHAPTER 1 INTRODUCTION Archean Nort h America is formed by the amalgamation of the Slave, Rae, Hearne, Nain, Superior, Wyoming, and Medicine Hat Block provinces. These areas of old, stable lithosphere are termed cratons, which represent proto continents which assembled over Earth history in to the continents. Ancestral North America, known as Laurentia, includes numerous crustal blocks with independent histories prior to this assembly. These blocks include the Slave, Rae and Hearne provinces of north central Canada, the Superior province, t he Nain province of eastern Canada, and the Wyoming province in Wyoming and Montana. In addition to the large variably exposed cratons, are smaller Archean tectonic elements. Many of these elements, such as the Sask Craton and Medicine Hat Block are conc ealed beneath younger sedimentary sequences. Laurentia also includes a series of Paleoproterozoic orogenic belts that represent the collisions between these cratons during continental assembly. The various provinces contain Early Proterozoic reactivation which appears to be related in trend and intensity to the orogenic belts and suture zones that frame them (Hoffman, 1988). The orogenic belts and suture zones range in age from ~2.0 Ga (Thelon orogeny suturing the Slave and Rae cratons) (van Breemen et a l., 1987; Tirrul and Grotzinger, 2010) to ~1.6 Ga (accretion of the Mazatzal terrane onto the Wyoming Province) (Karlstrom and Bowring, 1987). Hoffman (1988) concluded that Laurentia essentially accreted within the span of only 150 million years. The Wy oming Province is one of the oldest cratons in Laurentia (e.g., Wooden and Mueller, 1988; Baird et al., 1996; Frost et al., 1998; Henstock et al., 1998; Foster et al., 2006; Mueller and Frost, 2006). The craton is surrounded on all sides by

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20 Proterozoic su ture zones and orogenic belts resulting from continental margin accretion and collisions with other Archean cratons (e.g., Superior Wyoming collision at 1.71 1.77 Ga; Nelson et al., 1993; Dahl et al., 1999), or with Proterozoic terranes (e.g., Colorado pro vince at 1.78 Ga; Karlstrom and Houston, 1984; Chamberlain, 1998) ( Figure 1 1). Northwest of the Wyoming craton are the Great Falls tectonic zone (GFTZ) and the Medicine Hat block (MHB), which are largely defined by aeromagnetic, xenolith, and borehole da ta ( Figure 1 1) (Ross et al., 1991; Pilkington et al., 1992; Villeneuve et al., 1993; Baird et al., 1996; Gorman et al., 2002; Sims et al., 2004). The G reat F alls tectonic zone strikes northeast to southwest between the Archean Wyoming Province to the sou th, and the Archean Medicine Hat Block to the north ( Figure 1 1). Because the GFTZ is primarily covered by Phanerozoic sedimentary rocks, interpretations of the the G FTZ as a Paleoproterozoic suture, which has since been supported by geochronologic and geochemical data from Mueller et al. (2002) from gneisses of the Little Belt Mountains that revealed a subduction generated igneous arc signature in rocks formed during the time interval of 1.9 Ga to 1.8 Ga. Roberts et al. (2002) also supported the ocean subduction hypothesis based on 40 Ar/ 39 Ar analyses on biotite and 207 Pb/ 206 Pb step leach analyses on garnet from the Montana Metasedimentary Province (MMP). Garnets reve aled a metamorphic event between 1.82 and 1.79 Ga, followed by post tectonic cooling between 1.78 to 1.74 Ga (biotite) (Roberts et al., 2002), which was related to continental collision between Wyoming and the MHB along the GFTZ. Alternatively, Boerner et al. (1998) suggested that the GFTZ is an intra continental shear zone and that the MHB and Wyoming cratons are contiguous units. They

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21 propose that the subsurface Vulcan Structure represents the Hearne Wyoming suture and that based on potential field maps (Ross, 1991), the Deep Probe seismic experiment, Southern Alberta Refraction Experiment (SAREX) (Henstock et al, 1998), and electromagnetic studies (Boerner et al., 1998) the lithosphere is continuous throughout the GFTZ. This continuity would then link the geologic histories of the northern Wyoming Province and the MHB (e.g., Buhlman et al., 2000; Boerner et al., 1998). This dissertation addresses the nature of the Great Falls Tectonic zone crust, as well as characterizing GFTZ crust to evaluate possible links to the Wyoming Craton and/or Medicine Hat Block. Several specific objectives to address these issues include: 1) establish the age of crustal units, 2) identify the origins of the sampled crust in terms of plate tectonic settings, 3) examination o f the relationships between timing and nature be accomplished over three chapters, focused on the GFTZ segment in north central Montana. Chapters Two and Three uti lize crustal xenoliths entrained in magmas and ejected to the surface by Cenozoic magmatism throughout the Montana Alkali Province. These afford a unique opportunity to characterize the age and chemistry of the concealed crust of the GFTZ. Xenoliths are some of the only direct samples of the crystalline basement in central Montana. Chapter Two examines a crustal xenolith suite collected from the Grassrange diatremes ( Figure 1 2) of central Montana. Data generated by this study reveals major tectonic eve nts at c. 2.5 Ga and c. 1.75 Ga, related to subduction zone processes. Chapter Three focuses on a sample suite

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22 collected further north, from the Missouri Breaks diatremes and lavas in the Bearpaw and Highwood Mountains ( Figure 1 2). Abundant evidence of c. 2.5 2.6 Ga tectonic activity is preserved, as well as events at c. 1.75 1.89 Ga. Chapter Four examines the Precambrian rock exposed in the Little Rocky Mountains ( Figure 1 2). This domal uplift exposes some of the only crystalline basement of the MHB and provides a unique sampling opportunity. The rocks are dominantly Neoarchean, with both igneous ages and detrital zircon ages consistent with a Medicine Hat Block affinity. Through the study of basement exposures from the Little Belt and Little Roc ky Mountains, as well as through crustal xenoliths from seven locations across the Great Falls tectonic zone, this dissertation seeks to further define evolution of the GFTZ and the timing of ocean closure and collision between the Wyoming Province and the MHB as well as further characterize the age and composition of the crust beneath central Montana. Further, this study provides evidence on the extent of Archean basement in north central Montana, and suggests that the Archean basement therein is related to the MHB.

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23 Figure 1 1. (A) Paleoproterozoic Laurentia (after Hoffman, 1988; Foster et al., 2006; Davidson, 2008) and a l ocation map showing study area. (B) Generalized map of Precambrian basement provinces of southwestern Laurentia (after Ross et a l., 1991; Condie, 1992; Doughty et al., 1998; Vogl et al., 2004; Foster et al., 2006 2012 ). Exposures of basement in Laramide style uplifts are shown in the dark grey shaded areas.

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24 Figure 1 2 Generalized depiction of Cenozoic alkaline rock occurren ces in the Montana alkali province (after Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well defined and are based on aeromagnetic data from Sims et al., 2004.

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25 CHAPTER 2 GRASSRANGE XENOLITHS Introduction The Arc hean Wyoming Province is a geophysically and geochemically distinct entity within the North American continent surrounded by Paleoproterozoic orogenic belts (e.g., Wooden and Mueller, 1988; Baird et al., 1996; Frost et al., 1998; Henstock et al., 1998; Mue ller and Frost, 2006). The surrounding Proterozoic orogenic zones resulted from collisions with Paleoproterozoic terranes (e.g., Colorado province at 1.78 Ga; Karlstrom and Houston, 1984; Chamberlain, 1998) or other Archean cratons (e.g., Superior Wyoming collision at 1.71 to 1.77 Ga; Nelson et al., 1993; Dahl et al., 1999) ( Figure 2 1). There is significant uncertainty regarding the 700 km wide zone of Archean and Paleoproterozoic tectonic elements that separate the Wyoming Craton from the Hearne Provinc e. This zone includes features identified from geophysical surveys, borehole data (Villeneuve et al., 1993), and limited xenoliths (Davis et al., 1995; Gorman et al., 2002). These are from north to south in Figure 2 1: the E W trending Vulcan structure ( Eaton et al., 1999), the Archean Medicine Hat Block (e.g., Lopez, 1985; Mueller et al., 2002; Mueller et al., 2005; Foster et al., 2006). The Great Falls Tectonic Zone (GFTZ) strikes southwest to northeast between the Archean Wyoming Craton to the south and the Archean Medicine Hat block to the north and is largely covered by Phanerozoic sedimentary rocks. Consequently, interpretations of the origin and evolution of th (1985) identified the GFTZ and suggested that it formed a Paleoproterozoic suture zone between the Archean Medicine Hat and Wyoming cratons. Geochronologic and

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26 geochemical data (Mueller et al., 2002) provide addition al support for this hypothesis, revealing a subduction generated igneous arc signature in rocks formed during the time interval of 1.9 Ga to 1.8 Ga in the Little Belt Mountains. An alternative hypothesis views the Medicine Hat and Wyoming cratons as conti guous units separated by an intra continental shear zone (GFTZ) reactivated by involvement with the Paleoproterozoic Trans Hudson Orogeny (THO), with the Hearne Wyoming suture proposed to be the subsurface Vulcan Structure (Ross, 1991; Henstock et al. 1998 ; Boerner et al., 1998). These geophysically based models propose a continuity of lithosphere across the GFTZ, thereby linking the histories of the northern Wyoming Province and the Medicine Hat block (MHB) (e.g., Buhlman et al., 2000; Boerner et al., 199 8). Crustal xenoliths were ejected to the surface by Cenozoic magmatism in the Montana Alkali Provence within the GFTZ from the Bearpaw Mountains, Missouri Breaks diatremes, and Grassrange diatremes ( Figure 2 2). These xenoliths provide samples that can be used to obtain both age and compositional information for the crystalline basement in the GFTZ. We report U/Pb ages and Hf isotopic data from zircon, as well as whole rock geochemical and isotopic data, from a suite of crustal xenoliths from the Eocen e Grassrange intrusions (~50 Ma) of central Montana. These data give insight into the range of rock compositions and ages in the Great Falls tectonic zone and provide strong evidence that the GFTZ is a suture zone between the MHB and Wyoming craton. Geolo gic Background Extensive Phanerozoic sedimentary cover conceals much of the Precambrian geology of the GFTZ. The Little Belt Mountains (LBM) of Central Montana are a basement cored foreland structure of the Rocky Mountains intruded by Tertiary stocks

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27 (Kle inkopf et al., 1972; Marvin et al., 1973) and contain the most extensive outcrops of Proterozoic rocks in the GFTZ ( Figure 2 2). The Little Belt Mountain exposures include Paleoproterozoic dioritic orthogneisses and minor magmatic paragneisses (Pirsson, 1 900; Weed, 1900; Vogl et al., 2003). The gneissic fabric predominantly strikes east west and dips variably from north to south (Holm and Schneider, 2002; Vogl et al., 2003). The dioritic gneisses and migmatites are intruded by amphibolite dikes, pegmatit es, and post tectonic leucogranite (Vogl et al., 2003). Mueller et al. (2002) obtained an U Pb zircon age on the Pinto diorite of 1.86 Ga; other gneissic units yield U Pb zircon ages ranging from 1.88 1.86 Ga (Mueller et al., 2002). Vogl et al. (2003) an d Foster et al. (2006) give U Pb zircon ages of other 1.79 1.86 Ga dioritic gneisses, amphibolites and granitoids from the L BM exposures. Dahl et al. (2000) present U Pb data for monazite from a metapelite in the L BM which yielded an age of 1.86 Ga. Hol m and Schneider (2002) presented nine 40 Ar/ 39 Ar ages from the L BM ; six biotite separates yielded ages ranging from 1.81 Ga to 1.77 Ga, and three hornblende samples yielded a range of 1.80 Ga to 1.78 Ga. These 40 Ar/ 39 Ar ages record cooling after the last t ectonothermal activity in the L BM part of the GFTZ. Results U Pb and Lu Hf data from igneous and detrital zircons in addition to whole rock geochemical and isotopic data from xenoliths collected from two Eocene diatremes in the Grassrange, Montana are pre sented below Sample locations are shown in Figure 2 2. Geochemical and isotopic data and the latitude and longitude of the two sample locations are summarized in Tables B 1 through B 4 (Appendix B ).

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28 Whole R ock Geochemistry Xenoliths were divided into three categories on the basis of hand sample and thin section petrography, and major element geochemistry: granitic orthogneiss, quartzite/sandstone, and schist (Table 2 1 and 2 2 ). Representative photomicrographs for each group are shown in Figure 2 3 Granitic gneisses range in texture from hypidiomorphic granular to moderately foliated. They all contain blocky grains of microcline (>4mm), quartz, and minor biotite. Plagioclase is often sericitized ( Figure 2 3 A, B); rare titanite and zircon grains ar e also present. The granitoid samples have silica contents ranging from 73 to 77 wt. %. Sample MX 11 has a silica content of 80 wt. % suggesting silicification, because its trace element ratios are consistent with those of the lower silica samples (see be low). Collectively, the granitic samples cluster along the metaluminous peraluminous boundary ( Figure 2 4 ) when plotted according to the alumina saturation index of Shand (1943). The second and third groups consist of five biotite schists and three qu artzites/sandstones respectively. Schists are characterized by equigranular quartz, plagioclase, and K feldspars in all of the samples (ranging from 25 30% each) with clots of muscovite inter grown with sillimanite. Biotite defines a foliation along wi th accessory muscovite ( Figure 2 3 C, D). The protolith(s) of these schists may be sedimentary, volcaniclastic, and/or volcanic based on major element compositions (60 to 82% SiO 2 ). Abundances of Al 2 O 3 are relatively low (9.1 to 13.6 wt. %), but alkali co ntents are sufficiently high that their normative mineralogy contains 2 to 9% corundum. The third group consists of three un foliated biotite quartzites or sandstones. Quartz grains range from sub to well rounded, and some show clear euhedral authogenic overgrowths ( Figure 2 3 E, F). Visual estimates of other phases found within

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29 the quartzites/sandstones include plentiful biotite, oxides, and trace amounts of muscovite and feldspars. The metamorphic, grade of the quartzites is therefore lower than the s chists. Petrographic discrimination between sedimentary, volcaniclastic, and/or volcanic origins for the schists is difficult; however, variations of Ca, Mg, and Al concentrations provide some insight into their origins. Using the P 2 O 5 /TiO 2 versus MgO/Ca O plot of Werner (1987) to distinguish magmatic from sedimentary rocks indicates a likely sedimentary origin for most samples ( Figure 2 5 ). Shaw (1972) used a discriminant function with the equation: DF3 = 10.44 0.21SiO 2 0.32Fe 2 O 3 t 0.98MgO + 0.55CaO + 1.46Na 2 O + 0.54K 2 O to distinguish igneous versus sedimentary protoliths; igneous rocks have a positive DF3 value, and sedimentary rocks have a negative value ( Figure 2 6 ). Among the schistose samples, one lies within the igneous field, while the remain ing four are within the sedimentary field of Werner (1987) ( Figure 2 5 ). In the Shaw (1972) diagram ( Figure 2 6 ), one sample also plots within the igneous field and the remaining four within the sedimentary field. Although most of the schistose samples h ave compositional characteristics that indicate a sedimentary protolith, these discriminants are not sufficiently accurate to exclude a volcaniclastic component. Trace element geochemistry (Table 2 1 ) of the granitoid orthogneiss xenoliths is summarized in Figure 2 7 A, normalized relative to primitive mantle values of McDonough and Sun (1995). The trace element patterns of the granitoid samples are compared to the 5 schist samples in Figure 2 7 B. The plots show general enrichments in fluid mobile incomp atible elements, such as Rb, Ba, and Pb, up to 600 times the primitive mantle values. This relative enrichment is paired with minimal enrichment in

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30 fluid immobile trace elements, including heavy REE ( Figure 2 8 ). The majority of samples show lower values of Eu, Sr, and Ti relative to elements of similar compatibility. One sample, MX 18, lacks a Eu anomaly; however, it shares the Sr and Ti anomalies of the other samples. Nb and Ta are depleted relative to the observed values for neighboring elements in a ll of the granitic samples when normalized to primitive mantle U Pb Geochronology I n Zircon Meta granitoid samples MX 06, MX 08, MX 09, MX 10, MX 11, and MX 18 were large enough for whole rock geochemistry and zircon separation (~10 to 20 cm diameter). In dividual zircon 207 Pb/ 206 Pb ages from 153 grains (<10% discordance) from these samples reveal a wide spread, from 1.72 Ga to 2.55 Ga, despite their relatively uniform elemental compositions (Table 2 2). Two samples, MX 08 and MX 18, give Paleoproterozoic crystallization ages of 1.73 0.01 Ga ( Figure 2 9 A) and 1.74 0.01 Ga ( Figure 2 9 B) respectively, without inherited zircons. Zircons from the remaining four granitic samples reveal two distinct populations of 207 Pb/ 206 Pb ages, which are separated by > 5 Younger ages (lower intercepts) range from 1.76 0.02 Ga to 1.87 0.02 Ga (Table 2 2 ) and represent 30% of all grains <10% discordant. Ages in the older population (upper intercepts ) range from 2.46 0.16 Ga to 2.53 0.01 Ga and represent 17% of the <10% discordant grains. The remaining data yield intermediate ages that lie along a mixing line (discordia) between the older and younger ages, and constitute more than half of the ana lyses for each sample ( Figure 2 10 to 2 1 3 ). Two groups of metasedimentary samples were processed for detrital zircon analysis. The biotite schist samples MX 13, HAL 1, HAL 2, HAL 5, and HAL 6 yielded zircons ranging from sub rounded to well rounded. I n aggregate, 207 Pb/ 206 Pb ages from

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31 59 zircons <10% discordant range from ~1.75 Ga to ~3.06 Ga (Table 2 3), with prominent age peaks at ~1.86 G a (42% of the grains, from 1.82 to 1.90 Ga) and ~2.69 G a (15% of the grains, from 2.63 to 2.71 Ga) ( Figure 2 14 ). The three quartzite/sandstone samples (MX 12, MX 16, MX 17) yielded ninety five zircons <10% discordant with ages ranging from ~1.02 Ga to ~2.85 Ga (Table 2 3). The most prominent age peaks are illustrated in Figure 2 1 4 and lie at ~1.87 Ga (51% of the total, from 1.81 1.92 Ga), ~2.60 Ga (7% of the total, from 2.57 to 2.61 Ga), and ~2.68 Ga (8% of the total, from 2.67 to 2.72 Ga). Figure 2 1 4 shows a probability distribution function comparing detrital zircon age spectra from the schists and the quartzi tes/sandstones; data from the granitic samples are provided for comparison. Hf Isotopes I n Zircon Zircons were chosen for Lu Hf analysis based on their U Pb age (and discordance), zonation patterns, and grain size. Granitoid xenoliths MX 06, MX 09, MX 10 and MX 11 yielded upper intercept crystallization U Pb ages of 2.46 Ga to 2.53 Ga. (T) ~2.5 Ga) was calculated for each zircon regardless of their apparent age in order to ev aluate the (2.5 Ga) for these samples ranged from 3.0 to 6.7 (Table 2 2). When zed grain (1.8 to 2.5 ranged from 2.3 to 20.2 (Table 2 2). These differences likely reflect variable input of juvenile versus evolved materials in the genesis of the parental magma over the course of zircon crystallization. As suggested by the reduced ranges of Hf 2.5 averages for each sample ranged from 1.1 to 4.3, with standard deviations of 1.8 to 0.8. The more coherent Hf isotopic

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32 compositions calculated at ~2.5 Ga suggest that the younger ages reflect Pb loss and that the Lu Hf system remained largely closed during metamorphism. These samples give a limited range of depleted mantle model ages (T DM ) from 2.85 Ga to 2.93 Ga (Table 2 2). For the Paleoproterozoic samples (MX 08 and MX 18) th e measured Hf isotopic compositions for all zircons were recalculated to the U Pb crystallizat ion ages of (1.7 Ga) for <10% discordant zircons from the samples range from 4.7 to 12.0 for MX 08 with an average of 10.0 and a standard deviation of 1.4. MX 18 ranges from 7.4 to 15.0 with an average of 11 .6 and a standard deviation of 2.1 (Table 2 2). Average Hf T (DM) model ages for samples MX 08 and MX 18 were calculated to be 2.51 Ga and 2.58 Ga respectively. As with the older grains, calculating the initial Hf compositions using the upper intercept age (T) ) yields more coherent estimates of initial Hf compositions than using the 207 Pb/ 206 Pb age (IA) ), i.e., the discordant ages represent Pb loss rather than extraneous grains. The Hf T DM ages represent minimum mantle separation ag es because Lu/Hf in zircon is invariably lower than in whole rocks (e.g., Griffin et al., 2002). Calculation of secondary or crustal residence ages, however, require knowledge of the Lu/Hf of source(s) and were not calculated because the data ultimately s uggest a mixing of crustal and mantle sources at ~1.7 Ga. Only 3 of the 5 schist samples yielded zircons large enough for U Pb and Lu Hf isotopic analysis. Initial Hf isotopic compositions were calculated using the 207 Pb/ 206 Pb of each individual zircon. MX 13, HAL 2, and HAL <10% discordant grains that ranged from 2.8 to 9.9, 6.8 to 11.4 and 9.9 to 11.1, respectively (Table 2 3). MX 13 has Hf T (DM) model ages ranging from 2.13 Ga to 2.80

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33 Ga while HAL 2 and HAL 5 have Hf T (DM) model ages ranging from 1.98 Ga to 2.87 Ga and 1.85 Ga to 3.15 Ga, respectively (Table 2 3). Hf isotopes were also measured in zircons from quartzite/sandstone samples MX 12, MX 16, and MX 17, which yielded a to 14.4, 3.7 to 11.7, and 6.8 to 9.5, respectively (Table 2 3). The calculated Hf T (DM) model ages range from 1.61 Ga to 3.03 Ga for MX 12, 2.13 Ga to 2.99 Ga for MX 16, and 1.95 Ga to 2.99 Ga for MX 17 (Table 2 3). Sm Nd Whole Rock Isotopes As a gro (0) from 26.5 to 34.4. Initial ratios were calculated using the best estimates of the individual crystallization ages ( 207 Pb/ 206 Pb ages 2.46 Ga to 2.53 Ga) shown in Fig ures 2 10 to 2 1 3 and values for sample s MX 08 and MX 18 were calculated using the 207 Pb/ 206 Pb ages shown in Figure 2 9 (2.5 Ga) of 3.2 to 1.2 (Table 2 4, Figure 2 15 (1.7 Ga) of 6.0 and 6.5 (Table 2 4, Figure 2 1 5 ). Depleted mantle mo del ages for granitoid samples MX 06, MX 09, MX 10, and MX 11 were calculated using the model of DePaolo (1981) and ranged from 2.50 Ga to 2.79 Ga ( Figure 2 1 6 ). Granitic samples MX 08 and MX 18 yielded depleted mantle model ages (calculated at ~1.7 Ga; D ePaolo, 1981) of 2.47 Ga and 2.51 Ga (Table 2 4) respectively, overlapping with the ~2.5 Ga granitoids. The initial orthogneisses from the Little Belt Mountains (Mueller et al., 2002) and with an estimated Mueller et al., 1993; Frost, 199 (2.5 Ga) values are intermediate between 2.5 Ga depleted mantle and typical northern Wyoming province crust, suggesting involvement of both crustal and mantle sources. The five schist samples show a (0) values from 24.1 to 24.8, and depleted

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34 mantle model ages ranging from 2.35 Ga to 2.57 Ga (Table 2 4). The three quartzite (0) values from 13.7 to 23.8, and depleted mantle model ages ranging from 1.84 Ga to 2.42 Ga (Table 2 4). Discussion Origins O f T he Granitic Xenoliths Zircon U Pb age data show that the granitic xenoliths fall into two distinct ranges in terms of their crystallization ages, 2.46 Ga to 2.53 Ga for four of the six meta granitoids, and youn ger ages of 1.73 Ga and 1.74 Ga for the other two. The coherence of initial Hf isotopic data calculated at the proposed crystallization ages suggest these ages are robust and that the younger ages in each sample reflect loss of radiogenic Pb much more so than formation of new zircon with Hf from a higher Lu/Hf source (e.g., whole rock). The spread of sample data through the metaluminous and peraluminous fields ( Figure 2 4 ) is not particularly diagnostic, but is similar to granitic suites from continental arc and continental collision granitoids (e.g. Mainar and Piccoli, 1989; Rogers and Hawkesworth, 1989; Chappell and White, 2001; Villaseca et al., 2012). These xenoliths share relatively similar normalized trace element patterns that are consistent with t hose found in typical modern convergent margin igneous rocks ( Figure 2 7 A, Pearce, 1983). The ~2.5 Ga granitoid orthogneiss samples have negative Eu and Sr anomalies ( Figure 2 8 ), which typically represent fractional crystallization of plagioclase during their petrogenesis or residual plagioclase in the source. Figure 2 8 shows REE plots of the granitic samples normalized to chondritic values (McDonough and Sun, 1995). Depletion in the HREE is evident, indicating equilibrium with garnet and/or amphibole pyroxene bearing residue. Lastly, there are enrichments in fluid mobile elements such as Pb and Ba, which are commonly related to fluxing water from

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35 subducted materials in arc settings (Pearce, 1983; Thompson et al., 1984). Samples MX 08 and MX 18 yielde d Paleoproterozoic ages of 1.73 Ga and 1.74 Ga and represent magmas formed and crystallized at that time. Elemental abundances follow the same general trace element pattern observed in the older granitoid samples; however, the overall abundance of these e lements is lower ( Fig s 2 7 and 2 8 ). The evolved Nd and Hf isotopic signatures of the granitoids suggest two different histories for the ~1.7 Ga and ~2.5 Ga granitoids. Figure 2 15 illustrates the variation of sing the apparent 207 Pb/ 206 Pb age of each (2.5) (1.7) for the granitic samples. The older samples (MX 06, MX 09, MX 10, MX 11) crystallized at ~2.5 Ga, but have mea n Hf T DM for their zircons from 2.85 Ga to 2.93 Ga and (2.5) values from 6.7 to 3.0 ( Figure 2 1 7 ) The range of initial Hf isotopic compositions is largely mirrored in the (2.5) values and collectively suggest involvement of the depleted mantle and an older enriched component such as the northern Wyoming craton lithosphere ( Figure 2 1 5 Mueller et al., 2010; Mirnejad and Bell, 2006; Mirnejad and Bell, 2008). The mechanism of interaction with the older reservoir may have occurred by direct contact w ith ~2.8 Ga crust, or by incorporation of ~2.8 Ga detritus from the Wyoming craton that was being subducted (e.g., Mueller et al., 2010) and mixed into the sub lithospheric mantle wedge. The Paleoproterozoic samples (~1.7 Ga, MX 08 and MX 18) give mean H f T DM ages for their zircons of 2.53 Ga and 2.58 Ga respectively (Table 2 2 ), which constitute greater discrepancies between crystallization and model ages than observed for the older rocks, suggesting a stronger influence of older crust in their genesis. Shown on

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36 Figure 2 1 7 a line with a slope approximating the bulk continental crust evolution ( 176 Lu/ 177 Hf = 0.0115; Rudnick and Gao 2003) is plotted to coincide with the lower part (1.7) values for the c. 1.7 Ga samples cluster above the evolution line. Using the average of the Lu/Hf of the four xenoliths ( 176 Lu/ 17 7 Hf = 0.016) does not make an appreciable difference over this span of time. This relationship suggests that the ~1.7 granitic samples were not derived solely from partial melts of the ~2.5 Ga granites. The older component in these samples may be related to pre 1.8 Ga lithosphere, e.g., new lithosphere formed during the ~2.5 Ga event, or the older lithosphere that influenced the ~2.5 Ga magmas. The (1.7) values for zircons from the ~1.7 granites range from 4.7 to 15.0, suggesting mixing between a DM (depleted mantle) like component and an evolved crustal component(s), which likely include the ~2.5 granitoids. The (1.7) values calculated are 6.0 and 6.5, and plot below data for the Little Belt Mountains ( Figure 2 1 5 ), and seem to indicate a mixture between a DM like source and an older crustal component. However, it is important to note that these Paleoproterozoic samples are not the produc t of wholesale melting of the ~2.5 Ga granitic material based on their bulk compositions as well as Sm Nd and Lu Hf systematics ( Figure 2 1 7 ). They clearly contain portions of new crust likely formed during the ~1.8 1.9 Ga Little Belt subduction event (Mu eller et al., 2002; Vogl et al., 2003; Foster et al., 2006, 2012). Origins O f T he Schistose and Quartzite Xenoliths The biotite schist xenoliths are difficult to constrain in terms of origin because only a maximum age of deposition can be determined. The Al 2 O 3 to alkali ratio is sufficiently high in these rocks to produce normative corundum, which may be attributed to clays within the protolith, i.e., a possible indicator of a sedimentary component and

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37 consistent with the large amount of intergrown sillima nite and muscovite. These schistose xenoliths share relatively similar trace element patterns ( Figure 2 7 B) with the granitoid xenoliths ( Figure 2 7 A) and are consistent with those found in typical modern convergent margins (Pearce, 1984; Rogers and Hawke sworth, 1989; Taylor and McLennan, 1995; Chappell and White, 2001). As discussed below, the dominant zircon age peak within the schists (1.81 Ga to 1.90 Ga) corresponds to the time of Little Belt igneous activity, which suggests that the samples may conta in re worked volcaniclastic material from the Little Belt arc. Zircon 207 Pb/ 206 Pb ages from the quartzite and schist zircons that cluster between 1.81 Ga and 1.90 Ga ( Figure 2 1 4 ranging from 9.9 to 14.4 (Table 2 3). The 1.81 Ga to 1.90 Ga range of values for the Mountains, which range from values suggests that both juvenile (mant le like), and evolved sources contributed to the (0) for the schist samples is very small, from 24.8 to 24.1, and the Nd depleted mantle model ages range only from 2.34 Ga to 2.57 Ga (Table 2 4). Both the Hf in zircon and Nd in whole rock systematics indicate a mixture of sources in the provenance of these samples. The quartzites/sandstones share a similar Paleoproterozoic and Archean detrital zircon age distribution to the schistose samples ( Figure 2 1 4 ). They are not identical, however, because the minimum age of the youngest observed grains are Mesoproterozoic. Sample MX 16, contains detrital zircons with ages of ~1.0 Ga (3 grains) and MX 12 contains one ~1.3 Ga detrital zircon, while the youngest detrital grains in the schists are c. 1.75 Ga or older. The lack of a metamorphic overprint on

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38 these samples and presence of a small population of Mesoproterozoic zircon suggests that they are from part of the Phanerozoic cover sequence. The zircon popu lation is consistent with that in the Cambrian Flathead sandstone (Sears and Link, 2007; Mueller et al., 2008b). Implications F or G reat F alls T ectonic Z one Evolution The geochemistry and geochronology of the Grassrange xenolith suite and similar data from the Little Belt Mountains (Mueller et al., 2002; Vogl et al., 2003; Weiss et al., 2009) illustrate the heterogeneous nature of the central part of the GFTZ and its relationship to surrounding terranes. These relationships are shown in Figure 2 1 8 in whic h published Archean to earliest Proterozoic U Pb ages of zircon from the GFTZ are compared to ages from the northern Wyoming craton and the limited ages reported for the MHB. Ages were limited to those older than 2.4 Ga in order to make the most direct co mparisons of pre GFTZ data. Crystallization ages of the granitic samples (indicated by a star) as well as the detrital suites of zircons from the schist and quartzite samples are also plotted in Figure 2 18 Among the quartzite/sandstone and schist sampl es, those samples that contain a prominent age peak are indicated by a darker bar within the gray bar showing the range of zircon ages. The dominance of ~2.5 2.6 Ga ages suggests that they share a closer affinity to the MHB, which is characterized by Ne oarchean ages, than to the northern Wyoming Province, which is dominantly Mesoarchean ( Figure 2 1 8 ). The older (~2.7 to 2.8 Ga and ~3.0 Ga) Archean ages in the detrital samples are not distinctive, i.e., comparable to published data from both the MHB (Vil leneuve et al., 1993; Davis et al., 1995) and the northern Wyoming Province (e.g., Mueller et al. 2010). The two granitic samples interpreted to have formed 1.73 Ga to 1.74 Ga ago (MX 08 and MX 18) were most likely melts that crystallized during the

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39 final stage of the continent continent collision between the MHB and northern Wyoming craton, or post collision extension and underplating, based on their negative initial Hf and Nd compositions, i.e., considerably more negative than the values for samples of c omparable age from the Little Belt Mountains (Weiss et al., 2009). Consequently, the ages of the main Archean detrital zircon age peaks for the schists (2.63 Ga to 2.71 Ga) and the quartzites (2.57 Ga to 2.61 Ga) appear to be derived from Medicine Hat Blo ck or southern Hearne Province. Although limited in geographic distribution and age, the immobile trace element abundances for both Archean and Proterozoic meta granitoids support the proposition that the protoliths of both formed in a volcanic arc or sy n collisional setting or represent reworking of material initially formed in a subduction setting ( Figure 2 1 9 ; Pearce et al (1984)). In particular, these samples preserve the relative depletion in HFSEs ( Figure 2 7 A) characteristic of subduction zone mag matism (e.g., Thompson et al., 1984; Pearce et al., 1984) and reported for the Little Belt gneisses (Mueller et al., 2002). Trace element data from the Paleoproterozoic samples are easier to interpret because of the similarity in age to the Paleoproterozo ic gneisses in the Little Belt Mountains. These data are consistent with models proposed by Mueller et al. (2002, 2005) that the GFTZ represented a closing ocean basin and terminal continent continent collision between the Wyoming Craton and Medicine Hat block. The primarily subchondritic (1.7) (1.7) (whole rock) values for the granitic Paleoproterozoic samples ( Figure 2 1 7 ; Figure 2 1 5 ) both indicate substantial mixing between a DM like component and an evolved Archean crustal compone nt to a greater extent than evident in the slightly older (~1.86 Ga) igneous rocks in the Little Belt Mountains. This suggests

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40 that the ~1.7 Ga granitoids interacted with a significant amount of evolved material during their petrogenesis, a process consis tent with formation in a continental collision zone in which older crust was thickened and melted. The origin of the ~2.5 Ga gneisses is more difficult to constrain because they have no counterparts in either the Little Belt Mountains or in the Archean gn eisses exposed in the Beartooth Mountains or in southwestern Montana (see below). The U Pb ages from the schist and quartzite/sandstone samples are overlaps with the range for the Little Belt Mountains ( Figure 2 1 7 ). In this plot of zircon 207 Pb/ 206 quartzite samples appear to reflect a mixture of depleted mantle (defined by the upper dotte d line) and crustal sources (defined by the lower dashed line) in their parent rocks. crustal evolution line with a Lu/Hf ratio of 0.08 (Rudnick and Gao, 2003) in dicates more input from juvenile material and less from a more evolved (older) source for these Paleoproterozoic zircons. Mueller and Wooden (2012), for example, showed initial Hf values as low as 9 at 3.4 Ga for the northern Wyoming Province. In contr ast, the ~1.7 Ga granites show more negative Hf values at their time of crystallization, which suggests more intimate interaction with older reservoirs. Together, the initial Hf data from both igneous and metasedimentary rocks may be best viewed as a co ntinuum recording the transition from magma generation in a subduction setting where relatively primitive oceanic lithosphere was being subducted to a final continent continent

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41 collisional setting in which magmas received more input from older sources (e.g ., Mueller et al., 2011). To better understand the role of the GFTZ in the overall collisional history of Laurentia, it is important to note that collision of MHB and Wyoming coincided with other circum Wyoming province collisions (Mueller et al., 2011; Fi gure 2 2 0 ). For example, the Wyoming Colorado collision occurred between ~1.66 1.80 Ga (Bickford and Boardman, 1984; Sims and Peterman, 1986; Karlstrom and Bowring, 1988; Premo and Van Schmus, 1989; Van Schmus et al., 1993; Chamberlain, 1998; Selverstone et al., 2000; Hill and Bickford, 2001; Hill, 2004), and the Wyoming Superior collision to form the southern part of the Trans Hudson orogen occurred at 1.71 1.77 Ga (Karlstrom and Houston, 1984; Nelson et al., 1993; Resor et al., 1996; Dahl et al., 1999). The data presented here support the suggestion of Mueller et al. (2005) that the amalgamation of the Wyoming craton to the Superior craton and the Medicine Hat Craton was in part simultaneous, which places significant limits on plate tectonic relations du ring this very rapid period of continental growth (e.g., Hoffman, 1989; Mueller et al., 2005). Conclusions Geochronology and geochemistry of crustal xenoliths from the Grassrange provide new insight into the complex history of the crust in the GFTZ, incl uding: 1) Both Archean and Proterozoic granitic xenoliths show HFSE depletions and LIL enrichments characteristic of formation in a subduction modified environment. The limited enrichment of HREE relative to primitive mantle values reflects the source mi neralogy, likely lower crustal materials where garnet is residual after melt extraction. 2) U Pb data (<10% discordant) show distinct intervals (~1.7 and ~2.5 Ga) that are coincident with documented tectonothermal events in the MHB, suggesting that the b uried crust

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42 sampled by the Grassrange diatremes represents reworked MHB material in addition to juvenile Paleoproterozoic arc material similar to that in the Little Belt Mountains. 3) Zircon Hf and whole rock Nd isotopic data for the ~2.5 Ga granitic samp les indicate that juvenile and reworked crustal material mixed to varying degrees during crustal formation at ~2.5 Ga. These data, along with initial Hf isotope ratios for Archean detrital zircons, provide no evidence for crust as old as that which charact erizes the northern Wyoming Province (2.8 3.5 Ga). 4) The Hf and Nd isotopic data for the ~1.7 Ga granitic samples indicate a mixture of older crust (e.g., the ~2.5 Ga crust in the older xenoliths) and a more juvenile source. This is in contrast to expos ures of 1.8 1.9 Ga igneous material in the LBM, which involve far higher proportions of juvenile material, hypothesized to represent the arc formed during closure of the Little Belt ocean (Mueller et al., 2002). 5) The G reat F alls tectonic zone was initi ated prior to ~1.9 Ga as a convergent boundary related to the closing of an ocean basin between the Wyoming craton and the MHB that also produced the igneous suite of the Little Belt arc. The G reat F alls tectonic zone then evolved into a more transpressio nal boundary by ca. 1.77 as Wyoming moved east towards final collision with the Superior and Hearne provinces as well as the MHB (Dahl et al., 1999; Mueller et al., 2000; Mueller et al., 2002; Mueller et al., 2005). The compositions and ages of these xen oliths provide additional evidence for the evolution of the Great Falls tectonic zone via the closure of a Paleoproterozoic ocean and the addition of a significant volume of juvenile crust prior to continent continent collision. These data also suggest th at the arc was built on crust of the MHB rather than that of the Wyoming Province and that the arc evolved during the period of rapid coalescence of cratons that formed Laurentia in the Paleoproterozoic.

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43 Table 2 1 Major, trace and rare earth element da ta of granitoid xenolith s from the Grassrang e Major elements in wt. %, trace and rare earth elements in ppm. Sample MX 06 MX 08 MX 09 MX 10 MX 11 MX 18 wt. % Granitoids SiO 2 75 74 7 5 74 80 7 7 TiO 2 0.14 0.06 0.24 0.17 0.14 0.05 Al 2 O 3 1 3 1 4 1 4 1 3 1 1 13 Fe 2 O 3 0.33 0.47 1.04 0.94 0.51 0.70 MnO 0.01 0.01 0.01 0.01 0.01 0.01 MgO 0.05 0.13 0.24 0.25 0.30 0.06 CaO 0.71 0.62 0.92 0.63 0.23 0.97 Na 2 O 2. 9 3.8 3. 7 3.0 2.3 3. 9 K 2 O 6. 4 4. 4 5.0 6.5 6.1 4. 7 P 2 O 5 0.07 0.05 0.09 0.06 0.10 0.04 LOI 0. 84 1. 3 0.61 0.73 0.31 0.56 Total 99.20 99.40 100.25 99.54 100.43 100.92 ppm Li 5.1 6.9 5.9 4.2 7.9 4.2 Sc 1.5 2.1 1.6 1.4 1.6 1.2 Ti 841 47 1748 1073 913 51 V 14 12 16 18 14 12 Cr 4.1 4.3 7.5 6.7 4.6 6.0 Co 35 20 2.8 3.0 1.1 2.6 Ni 3.4 3.0 3 .8 4.1 3.6 3.5 Cu 4.3 24 6.2 3.5 2.9 4.3 Zn 20 25 29 23 24 24 Ga 15 23 14 14 13 16 Rb 170 164 130 161 207 129 Sr 103 67 138 91 70 105 Y 16 4.6 19 9.2 9.3 3.3 Zr 34 39 29 33 13 57 Nb 22 16 30 21 25 8.8 Cs 0.47 1.7 0.40 1.0 1.9 2.6 Ba 900 4 30 902 804 736 636 La 73 23 104 67 57 23 Ce 146 42 203 130 106 41 Pr 15 4.3 21 13 11 4.3 Nd 48 14 65 41 36 14 Sm 8.3 2.4 10 6.2 5.8 2.4 Eu 0.83 0.41 1.1 0.85 0.79 0.68 Gd 5.8 1.5 7.0 4.1 4.1 1.6 Tb 0.77 0.18 0.88 0.49 0.51 0.20 Dy 3.6 0.8 5 4.0 2.2 2.3 0.86 Ho 0.57 0.15 0.67 0.34 0.34 0.14 Er 1.5 0.45 1.6 0.84 0.81 0.35 Tm 0.21 0.07 0.18 0.10 0.09 0.05 Yb 1.4 0.55 0.92 0.60 0.44 0.34 Lu 0.19 0.09 0.10 0.08 0.05 0.06

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44 Table 2 1. Continued. Sample MX 06 MX 08 MX 09 MX 10 MX 11 MX 1 8 ppm Granitoids Hf 1.2 1.8 0.95 1.1 0.45 2.4 Ta 3.0 2.9 1.0 0.91 1.5 0.68 Pb 36 33 49 29 29 42 Th 42 6.6 77 40 34 19 U 2.0 1.7 6.8 2.5 1.7 2.9

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45 Table 2 2 Major, trace and rare earth element data of schist and quartzite xenolith s f rom the Grassrange Major elements in wt. %, trace and rare earth elements in ppm. Sample MX 13 HAL 1 HAL 2 HAL 5 HAL 6 MX 12 MX 16 MX 17 wt. % Schists Quartzites SiO 2 78 80 76 82 65 84 83 80 TiO 2 0.51 0.51 0.38 0.40 0.69 0.40 0.14 0.37 Al 2 O 3 11 10 13 9.2 17 9.1 2.3 9.9 Fe 2 O 3 3.2 3.4 2.8 2.4 4.8 0.76 2.3 2.6 MnO 0.03 0.03 0.02 0.02 0.05 0.01 0.07 0.02 MgO 0.97 1.0 0.91 0.68 1.7 0.32 0.46 0.86 CaO 0.46 0.35 0.29 0.66 0.76 0.47 4.7 0.61 Na 2 O 1.5 0.76 0.96 1.9 2.16 1.06 0.04 2.0 K 2 O 3.9 2.7 4.0 2.7 6.22 1.98 0.36 2.9 P 2 O 5 0.12 0.15 0.14 0.13 0.15 0.13 0.28 0.14 LOI 0.87 1.0 1.2 1.2 1.9 1.6 5.1 0.74 Total 100.26 100.23 99.77 100.94 100.27 100.10 99.10 100.37 ppm Li 25 42 39 19 42 16 24 21 Sc 7.4 8.6 7.5 6.6 14.69 4.93 1.54 5.73 Ti 4236 2841 2185 2285 4586 2970 518 2851 V 48 51 47 42 90 37 48 40 Cr 46 47 41 37 76 36 21 34 Co 6.4 6.9 5.7 4.8 11 38 8.1 12 Ni 18 18 15 12 35 5.1 13 15 Cu 7.3 5.7 11 14 6.8 8.7 16 7.8 Zn 57 58 50 40 92 24 33 57 Ga 12 13 1 5 9.6 22 11 2.6 10 Rb 148 136 115 108 230 74 12 120 Sr 119 56 80 100 171 66 84 99 Y 8.7 15 12 12 13 9.8 8.3 10 Zr 167 263 153 281 216 135 36 107 Nb 16 12 11 13 24 14 9.1 11 Cs 7.7 7.0 5.4 4.9 13 1.8 0.45 5.8 Ba 694 430 818 503 1193 330 1207 568

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46 Table 2 2. Continued. Sample MX 13 HAL 1 HAL 2 HAL 5 HAL 6 MX 12 MX 16 MX 17 ppm Schists Quartzites La 37 41 36 39 47 31 10 31 Ce 77 84 79 79 96 64 19 64 Pr 8.6 9.5 8.6 8.5 10 7.0 2.4 7.2 Nd 30 33 31 30 36 24 9.0 25 Sm 5.4 6.3 5.7 5.3 6.5 4.3 1.8 4.7 Eu 1.1 0.94 1.1 0.96 1.4 0.73 0.45 0.93 Gd 3.9 5.1 4.6 4.3 4.9 3.1 1.7 3.8 Tb 0.50 0.67 0.59 0.55 0.63 0.42 0.24 0.51 Dy 2.2 3.3 2.8 2.7 2.9 2.1 1.3 2.4 Ho 0.34 0.56 0.46 0.46 0.49 0.37 0.26 0.38 Er 0.82 1. 5 1.2 1.2 1.2 0.99 0.74 0.92 Tm 0.10 0.19 0.15 0.16 0.16 0.14 0.10 0.12 Yb 0.66 1.3 0.95 1.1 1.1 0.94 0.68 0.75 Lu 0.10 0.20 0.15 0.17 0.17 0.14 0.10 0.11 Hf 4.9 7.5 4.4 7.9 6.2 4.0 1.0 3.2 Ta 0.88 0.82 0.67 0.88 1.6 2.9 0.25 0.96 Pb 22 1 4 26 20 32 8.2 4.4 21 Th 17 17 17 14 17 10 2.3 13 U 4.2 4.1 3.8 3.5 3.8 2.9 2.4 3.2

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47 Table 2 3 Granitoid xenoliths, Grassrange, zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf isotope data reported (Ma). Sample Young Error Old Error (IA) (IA) (T) (T) (T) Std. Hf Model Age Std. Number U Pb Age U Pb Age Highest Lowest Highest Lowest Average Dev. (DM) a Dev. MX 06 1842 110 2524 25 2.3 19.9 3.0 3.6 1.1 1.8 2849 70 MX 08 1731 5 N/A N/A 0.5 11.8 4.7 12.0 10.0 1.4 2506 54 MX 09 1765 37 2519 15 0.6 19.3 0.4 5.5 2.0 1.8 2878 73 MX 10 1828 16 2528 8 3.0 19.5 1.4 3.8 2.9 0.8 2923 36 MX 11 1818 21 2456 160 4.6 20.2 0.6 6.7 4.3 1.7 2932 74 MX 18 1736 9 N/A N/A 7.5 15.0 7.6 15.0 11.6 2.1 258 1 94 (IA) Zircons reduced to individual U Pb age. *zircons reduced to old U Pb age, if "Old U Pb Age" column is "N/A", sample reduced to Young U Pb Age. a Depleted mantle model ages were calculated using the model of Mueller et al. (2008).

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48 Table 2 4 Sch ist and quartzite xenoliths, Grassrange, zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf isotope data reported (Ma). Sample Age Age Age (IA) (IA) Hf T (DM) a Hf T (DM) a Number Range Population Population Highest Lowest Highest Lowest Schist MX 13 1848 2409 1897 N/A 2.8 9.9 2804 2129 HAL 1 1809 2531 1864 N/A N/A N/A N/A N/A HAL 2 1818 2666 1899 2628 6.8 11.4 2874 1982 HAL 5 1839 3054 1866 2679 9.9 11.1 3148 1850 HAL 6 1751 1871 1781 N/A N/A N/A N/A N/A Quartzite MX 12 1282 2684 1871 N/A 5.8 14.4 3027 1610 MX 16 1018 2851 1847 N/A 3.7 11.7 2985 2126 MX 17 1813 2690 1864 2680 6.8 9.5 2889 1950 (IA) Zircons reduced to individual U Pb age. a Depleted Mantle model ages were calculated using the model of Muell er et al. (2008).

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49 Table 2 5 LA ICP MS Nd isotope data reported (Ma). Sample Sm Nd 143 Nd/ 144 Nd (0) (T) Nd Model Age Number (ppm) (ppm) (DM) a Granitoids MX 06 8.27 47.83 0.51120 28.0 1.9 2592 MX 08 2.41 14.00 0.51128 26.5 6.0 2465 MX 09 10.20 65.44 0.51088 34.4 1.2 2791 MX 10 6.18 41.24 0.51104 31.2 3.2 2501 MX 11 5.82 36.28 0.51108 30.4 1.0 2584 MX 18 2.37 13.61 0.51126 26.8 6.5 2513 Schists MX 13 5.36 29.79 0.51139 24.4 N/A 2422 HAL 1 6.28 33.4 4 0.51141 24.1 N/A 2511 HAL 2 6.28 33.44 0.51137 24.8 N/A 2574 HAL 5 5.71 31.16 0.51138 24.6 N/A 2484 HAL 6 5.27 30.12 0.51139 24.3 N/A 2345 Quartzites MX 12 4.25 23.98 0.51142 23.8 N/A 2341 MX 16 1.82 9.00 0.51194 13.7 N/A 1840 M X 17 4.74 25.46 0.51144 23.3 N/A 2421 *Granitic sample reduced to old age peak U Pb age (Table 2 3), if "Old U Pb Age" column is empty, sample reduced to "Young U Pb Age" column. a Depleted Mantle model ages were calculated using the model of DePaolo (1981).

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50 Figure 2 1 Generalized map of Precambrian basement provinces of southwestern Laurentia (after Ross et al., 1991; Condie, 1992; Vogl et al., 2004; Foster et al., 2006). Exposures of basement in Laramide style uplifts are shown in the dark grey shaded areas.

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51 Figure 2 2 Generalized depiction of Cenozoic alkaline rock occurrences in the Montana alkali province (after Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well define d and are based on aeromagnetic data from Sims et al., 2004.

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52 MX 08 (granitoid) (A) (B) HAL 1 (schist) (C) (D) MX 16 (sandstone) (E) (F) Figure 2 3 Representative exampl es of photomicrographs of (A) granitic sample MX 08 in plain polarized light (PPL); (B) granitic sample MX 08 in cross polarized light (XPL); (C) schist sample HAL 1 in PPL; (D) schist sample HAL 1 in XPL; (E) quartzite sample MX 16 in PPL; (F) quartzite s ample MX 16 in XPL. Scale bars are 1mm in size. 1mm 1mm 1mm 1mm 1mm 1m m

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53 Figure 2 4 Granitoid samples plotted on an alumina saturation index (Shand, 1943; modified by Frost et al., 2001). ~2.5 Ga granitoids (diamonds), ~1.7 Ga granitoids (squares).

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54 Figure 2 5 Plot of P 2 O 5 /TiO 2 vs. MgO/CaO (after Werner, 1987) proposed to discriminate magmatic vs. sedimentary protoliths for all meta igneous granitoids and schistose samples.

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55 Figure 2 6 Compositional discriminant diagram showing the discriminant function 3 (DF3) vs. SiO 2 for all granitoid and schistose samples (after Shaw, 1972). Rocks with positive DF3 values are interpreted to be of igneous origin, and rocks with a negative value are interpreted to have a sedimentary origin.

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56 Figure 2 7 Primitive mantle normal for ( A) granitoid samples and ( B) schistose samples, using data in Table 2 1 and 2 2 Relative enrichment of large ion lithophile elements ( LILEs ) to high field strength elements ( HFSE ) suggests formation in an arc environment.

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57 Figure 2 8 Chondrite normalized rare earth element diagram (McDonough and Sun, 1995) for granitoid samples using data in Table 2 1.

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58 Figure 2 9 Concordia diagrams showing U Pb data for granitoid sam ples MX 08 (A) and MX standard error.

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59 Figure 2 10 (A) Concordia diagrams plotting composite U Pb data for granitoid sample MX 06. Each ellipse represents error. (B) Expanded view of lower end of concordia showing samples selected for regression. (C) Expanded view of upper end of concordia showing samples selected for regression.

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60 Figu re 2 11 (A) Concordia diagrams plotting composite U Pb data for granitoid sample MX error. (B) Expanded view of lower end of concordia showing samples selected for regression. (C) Expanded view of upper end of concordia showing samples selected for regression.

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61 Figure 2 12 (A) Concordia diagrams plotting composite U Pb data for granitoid sample MX 10. Each ellipse represents a single spot analysis and error. (B) Expanded view of lower end of concordia showing samples selected for regression. (C) Expanded view of upper end of concordia showing samples selected for regression.

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62 Figure 2 1 3 (A) Concordia di agrams plotting composite U Pb data for granitoid sample MX error. (B) Expanded view of lower end of concordia showing samples selected for regression. (C) Expanded view of upper end of concordia showing samples selected for regression.

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63 Figure 2 1 4 Probability density plots comparing the cumulative <10% discordant zircon analyses of (A) the 6 granitoid samples, 155 grains; (B) The 5 schist samples, 59 grains; (C) The 3 quartzite samples, 95 grains.

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64 Figure 2 1 5 Sm Nd evolution diagram showing the range for the ~1.7 granitoids (gray stars), ~2.5 granitoids (white stars), Little Belt mountains (white bar), and the generalized evolution of northern Wyoming province crust (Wooden and Mueller, 1988; Mueller et al., 1993; Frost, 1993). CHUR chondritic uniform reservoir (Faure and Mensing, 2005).

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65 Figure 2 16 Histogram showing Sm Nd depleted mantle model ages (Ga) of granitoid, schist, and quartzite whole rocks. Depleted Mantle model ages calculated using the model of DePaolo, 1981.

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66 Figure 2 17 Hf evolution diagram showing the range for the ~1.7 granitoids (circles), ~2.5 granitoids (diamonds), schistose (square), quartzite (triangle) and LBMs (black dashes) from ages 1.70 Ga to 1.89 Ga and from 2.40 Ga to 2.75 Ga. CCBulk bulk continental crust calculated using Rudnick and Gao (2003) (dashed line), CHUR chondritic uniform reservoir, BSE bulk silicate earth (Faure and Mensing, 2005) (solid line), DM depleted mantle (C hauvel and Blichert Toft, 2001)(dotted line).

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67 Figure 2 1 8 Published Archean to earliest Proterozoic U Pb ages of zircon from the northern Wyoming craton (Heimlich and Banks, 1968; Mueller et al., 1995; Kirkwood, 2000; Mogk et al., 2004; Frost and Fann ing, 2006; Mueller et al., 2010, Krogh et al., 2011) compared to published, ages for the MHB (Villeneuve et al., 1993; Davis et al., 1995). OC Owl Creek Mountains, BM Bighorn Mountains, BT Beartooth Mountains, TR Tobacco Root Mountains, SG Sweetg rass Hills, BH borehole. Wyoming light gray, Medicine Hat dark gray. Crystallization ages of granitoids from this study stars, from schists (detrital) medium gray with dashed lines, and quartzite (detrital) medium gray with solid lines. Promi nent age peaks from Figure 2 1 4 black bars.

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68 Figure 2 19 Trace element discrimination diagrams (after Pearce et al., 1984) for granitoid samples: (A) H eavy rare earth element Y ppm vs. high field strength element Nb ppm, (B) H eavy rare earth element Yb ppm vs. high field strength element Ta ppm. ~1.7 granitoids are shown in orange circles and ~2.5 granitoids are shown in blue diamonds.

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69 Figure 2 20 Comparison of published U Pb ages from the southern Trans Hudson Orogen (Karlstrom and Houston, 1 984; Nelson et al., 1993; Resor et al., 1996; Dahl et al., 1999) against the northern Trans Hudson Orogen (Lewry et al., 1987; Hoffman, 1988; Bickford et al., 1990; Gordon et al., 1990; Machado, 1990; Lewry et al., 1994; Lucas et al., 1996; Sun et al., 199 6; Machado et al., 1999; Mueller et al., 2002), Great Falls tectonic zone (this study), and Yavapai Central Plains (Bickford and Boardman, 1984; Sims and Peterman, 1986; Karlstrom and Bowring, 1988; Premo and Van Schmus, 1989; Van Schmus et al., 1993; Cham berlain, 1998; Selverstone et al., 2000; Hill and Bickford, 2001; Hill, 2004).

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70 CHAPTER 3 MISSOURI BREAKS XENOLITHS Introduction Laurentia is composed of a number of Archean cratons sutured together across several Proterozoic mobile belts. The Archean Wy oming Province is one of the oldest Laurentian cratons, and is surrounded on all sides by Paleoproterozoic orogenic belts and suture zones (e.g., Wooden and Mueller, 1988; Baird et al., 1996; Frost et al., 1998; Henstock et al., 1998; Foster et al., 2006; Mueller and Frost, 2006). The Proterozoic mobile belts resulted from continental margin accretion and collisions with other Archean cratons (e.g., Superior Wyoming collision at 1.71 1.77 Ga; Nelson et al., 1993; Dahl et al., 1999), or with Proterozoic ter ranes (e.g., Colorado province at 1.78 Ga; Karlstrom and Houston, 1984; Chamberlain, 1998) ( Figure 3 1). The Great Falls tectonic zone (GFTZ) is one such zone, proposed to have formed by convergence and collision between the Wyoming Craton and the Medicin e Hat Block (MHB) to the north ( Figure 3 1) (Ross et al., 1991; Pilkington et al., 1992; Baird et al., 1996; Mueller et al., 2002; Gorman et al., 2002; Sims et al., 2004). The G reat F alls tectonic zone was first described as a Paleoproterozoic suture zone (1985). Geochronologic and geochemical data by Mueller et al. (2002), taken from meta igneous gneisses in the Little Belt Mountains (LBM) reveal a subduction generated igneous arc s ignature in rocks formed during the interval of 1.9 Ga to 1.8 Ga. Further support came from Roberts et al. (2002) who obtained biotite 40 Ar/ 39 Ar and 207 Pb/ 206 Pb step leached garnet ages from the Montana meta sedimentary terrane (MMT, Figure 3 1) and inter preted them to reflect a metamorphic event between 1.82

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71 and 1.79 Ga, followed by post tectonic cooling between 1.78 and 1.74 Ga. Pressure temperature time paths (Harms et al., 2004) from three metamorphic suites from the Tobacco Root Mountains and metamor phosed mafic dikes and sills revealed crustal thickening, metamorphism, and partial melting occurring between 1.78 and 1.76 Ga. An alternative origin for the GFTZ was suggested by Boerner et al. (1998), who hypothesized that the GFTZ is an intra continent al shear zone, and that the MHB and Wyoming cratons are contiguous units. They proposed that the Hearne Wyoming suture is associated with the Vulcan magnetic structure (Ross, 1991; Henstock et al. 1998; Boerner et al., 1998; Figure 3 1) and that, based on geophysical models, there is a continuity of lithosphere throughout the GFTZ and northern Wyoming Province. This interpretation would then link the histories of the northern Wyoming Province and the MHB (e.g., Buhlman et al., 2000; Boerner et al., 1998). Crustal xenoliths entrained in Cenozoic magmas erupted in the Montana Alkali Province afford a unique opportunity to characterize the age and chemistry of the subsurface of the GFTZ. This paper expands on crustal xenolith data from the Grassrange diatr emes, Montana presented by Gifford et al. (in review) (Chapter 2). Whole rock major and trace element geochemistry, whole rock Sm Nd isotopes, and zircon U Pb geochronology and Hf isotopic compositions were obtained from 17 crustal xenoliths from 5 additi onal locations within the Montana Alkali Province ( Figure 3 2). Three xenoliths from the Big Slide diatreme (Missouri Breaks), 5 xenoliths from the Little Sand Creek diatreme (Bearpaw Mtns.), 4 xenoliths from the Robinson Ranch diatreme (Bearpaw Mtns.), 4 xenoliths from a location within the Bearpaw Mountain volcanics called Lloyd Divide, and 1 xenolith from the Highwood Mountains were included in this

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72 study ( Figure 3 2), which seeks to further define the timing of ocean closure and collision between the M HB and the Wyoming Province, as well as test models for the Paleoproterozoic evolution of the Great Falls tectonic zone. Geologic Background The Wyoming Craton and the largely concealed Medicine Hat block ( Figure 3 1) are separated by the Great Falls tec tonic zone (GFTZ), a region of Paleoproterozoic tectonic activity (c. 1.86 to 1.71 Ga) that likely records the closure of an ocean basin timing of tectonic activity in the GFTZ is of particular interest due to its overlap with the Laurentia forming collisions in the Trans Hudson orogen (c. 1.83 1.72 Ga; Bickford et al., 1990; Dahl et al., 1999). Documentation of the MHB is largely from aeromagnetic and seismic data, s upplemented by samples from a few deep boreholes (e.g., Ross, 2002). Geochronology of crystalline basement collected from cores penetrating MHB crust reveals a range of ages from 3.27 Ga to 2.65 Ga (Villeneuve et al., 1993). The northern boundary of the block is the aeromagnetically defined Vulcan low, interpreted by Ross (2002) to mark a Proterozoic collision between the Medicine Hat Block and the Hearne Craton. Basement rocks are exposed in three main areas of the northwestern Wyoming craton (Bighorn a nd Beartooth ranges of the Bighorn Beartooth magmatic zone, and the Madison, Tobacco Root, etc. of the Montana Meta sedimentary Terrane), uplifted by Laramide thick skinned thrusting. These uplifts are identified as sharing a common heritage primarily due to their similar Pb isotopic data (Wooden et al., 1988; Mogk et al., 1992; Chamberlain et al., 2003; Frost, 1993; Mueller and Frost, 2006). Meta igneous rocks of the Montana meta sedimentary province record ages from 3.3 Ga to 3.1 Ga

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73 (Mueller et al., 199 6). The Beartooth and Bighorn mountains contain abundant metamorphosed tonalite trondhjemite granodiorite (TTG) suites with protolith ages between 2.9 to 2.8 Ga (Frost et al., 2006a; Mueller et al., 2010). The southern Wyoming Craton contains igneous arc terranes formed and accreted at 2.68 Ga to 2.5 Ga (Frost et al., 2006b). Dike emplacement at ~2.1 Ga is recorded across the Wyoming craton, possibly related to a period of Paleoproterozoic rifting (Premo and Van Schmus, 1989; Cox et al., 2000; Mueller et al., 2004). The G reat F alls tectonic zone of North Central Montana hosts numerous occurrences of Cenozoic alkalic magmatism and diatreme emplacement. These include localities in the Bearpaw Mountains, Highwood Mountains, Sweetgrass Hills, Crazy Mountains and the Grassrange and Missouri Breaks diatreme swarms ( Figure 3 2). Many of these localities contain crustal and upper mantle xenoliths (Collerson et al., 1989; Hearn, 1989; Hearn et al., 1989; Joswiak, 1992; Carlson and Irving, 1994; Downes et al., 200 4; Bolhar et al., 2007; Facer et al., 2009; Blackburn et al., 2010, 2011). Crustal xenoliths described in this study are from the Highwood Mountains, Big Slide Diatreme, Robinson Ranch diatreme, Little Sand Creek diatreme, and Lloyd Divide locality; xenoli ths are carried by Eocene minettes (54 50 Ma) of the Bearpaw Mountain volcanic field (Marvin et al., 1980; Hearn, 1989; MacDonald et al., 1992). U Pb ages of zircons from xenoliths collected in the Little Sand Creek locality are from ca. 3.0 to 1.8 Ga (Bo lhar et al., 2007). Mafic granulite xenoliths from the Grassrange Hills yield ages at c. 1.8 Ga, and Sm Nd isochrons (garnet, clinopyroxene, and whole rock) ranging from 1.7 to 1.5 Ga (Davis et al., 1995), which may represent either metamorphosed Archean crust or an addition to the lower crust ca. 1.8 Ga. Barnhart et

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74 al. (2012) analyzed 5 granulite xenoliths from the Robinson Ranch diatreme ranging from mafic to felsic in nature, and dating was done on monazite grains using the U Th total Pb procedure and yielded three age populations: c. 1.3 Ga to c. 1.5 Ga; c. 1.7 Ga to c. 1.8 Ga; and c. 2.0 Ga to c. 2.1 Ga. Barnhart et al. (2012) interpreted these populations as representing an incremental assembly of the high velocity lower crustal layer suggested by Gorman et al. (2002) during the formation of the GFTZ. Sample Descriptions Big Slide Diatreme BSD10 04 is a medium to fine grained quartzofeldspathic meta leuco granitoid. Plagioclase occurs as annealed, rounded, and sutured grains 0.5 2 mm in diameter, and comprising ~40% of the sample. Quartz grains are anhedral, lobate, and up to 2 mm in diameter. Undulose extinction is present in many quartz grains. Trace occurrences of perthitic K feldspar, biotite, muscovite, and opaque phases are also present. BSD10 05 is a mafic meta granitoid ( Figure 3 4A). Biotite is the dominant phase, occurring both as large clots that appear pseudomorphic after pyroxene or amphibole, and finer grains scattered throughout the matrix. The matrix is dominated by fine grain plagioclase are present as well. BSD10 06 is a medium grained garnet granulite with abundant clinopyroxene. It exhibits granoblastic texture with individual grains ranging up to 2 mm in diam eter. Robinson Ranch Diatreme RRD10 05 is an amphibole biotite meta granitoid ( Figure 3 3B). Amphibole The biotite is unaltered and generally in contact with the amph ibole grains. Plagioclase

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75 and microcline are subhedral and blocky, with a wide range of grain sizes (<1 3 mm in diameter). Quartz is abundant, generally anhedral to subhedral, and shows a similar range of sizes to the feldspar grains. The sample shows fractures. Sample RRD10 09 is similar, but displays a finer grained texture and more extensive alteration in the form of larger biotite reaction rims around amphibole and sericitized feldspars. RRD10 13 is a garnet granulite with 2 3 mm subhedral poikil itic garnets, which contain abundant fractures and rounded quartz and feldspar inclusions. There are domains rich in fine grained blocky sericitized plagioclase and anhedral quartz. Annealing textures on some of the quartz is common as well. Trace amoun ts of muscovite and biotite are scattered throughout the matrix. The sample shows abundant fractures as well as high temperature hornblende and pyroxene veins. Evidence suggests that the sample underwent retrogressive metamorphism at amphibolite facies. RRD10 20 is a highly altered and highly retrogressively metamorphosed dioritic granulite. The sample contains large amounts of hornblende and biotite. There are sections which appear to have been recrystallized, and clots of biotite contain cores which might have been pyroxene. Fine grained recrystallized, elongate quartz grains and ribbons range throughout the sample. Little Sand Creek Diatreme LSC10 03 is a fine to medium grained, lineated, biotite and amphibole rich rock ( Figure 3 3C). Amphibole is approximately 1 2 mm long and subhedral, with some grains showing poikilitic texture and containing inclusions and/or alteration minerals. Plagioclase grains are up to 3 mm in diameter, blocky, and generally subhedral. Quartz occurs as finer subhedral grains, up to 0.3 mm in diameter and many grains display undulose extinction. Preferred orientation of amphibole grains defines a lineation, with

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76 a weak compositional layering between amphibole rich and quartzofeldspathic rich domains defining a fabric. There are large amounts of opaque minerals. LSC10 10 is a length, although the grain size varies in compositional bands through the sample. There is fracturing and alteration throughout differen t zones of the sample. Plagioclase occurs as subhedral and blocky grains up to 0.5 mm in diameter, with grain size varying depending on the layer. LSC10 13 is a mafic granulite gneiss. Biotite is present, but does not form a foliation within the sample. Minor amounts of anhedral plagioclase are scattered around the sample. Euhedral grains of amphibole and clinopyroxene are up to 2 mm, blocky, and prevalent throughout the sample. LSC10 11 is a highly altered mafic gneiss ( Figure 3 4B). The sample cont ains foliated clots of intergrown high temperature mafic minerals as well as highly altered amphibole and plagioclase. Chlorite and biotite are present throughout the sample along with abundant opaques. LSC10 12 is a highly altered and silicified meta gr anitoid ( Figure 3 4C). Potassium feldspar grains display tartan twinning, and twinned plagioclase and quartz are present. There is highly chloritized biotite as well as trace amounts of muscovite. The plagioclase grains are pervasively sericitized and c lay altered. Within the plagioclase grains, there is evidence of myrmekitic texture. Bearpaw Mountains A t Lloyd Divide Samples LD10 01 ( Figure 3 3A), LD10 07, and LD10 08 are garnet quartzofeldspathic gneisses. Compositional banding is defined by garne t rich and quartzofeldspathic layers. Garnet grains are up to 1 mm and are anhedral. Quartz and plagioclase is finer and anhedral, although some plagioclase grains are blocky. LD10

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77 11 appears to be 98% coarse quartz (>1 cm in diameter), with trace amoun ts of amphibole and pyroxene. Highwood Mountains Sample HX 1 is a biotite quartzofeldspathic gneiss. The sample contains biotite rich and K feldspar rich bands, which are several mm thick. The sample fabric is defined by several elements: a biotite fol iation; recrystallized, elongate quartz grains and ribbons; and plagioclase feldspar augen. Results U Pb and Lu Hf data from igneous and metamorphic zircons in addition to whole rock geochemical and isotopic data from xenoliths collected from three Eocen e diatremes in the Missouri Breaks, as well as one from the Bearpaw Mountain extrusives, and one xenolith from the Highwood Mountains are presented below. Sample locations are shown in Figure 3 2. Methods are described in Appendix A. Geochemical and iso topic data and the latitude and longitude of the sample locations are summarized in Tables B 4 through B 6 (Appendix B). Whole Rock Geochemistry All of the samples from this study (except LD10 11) are meta igneous. Representative photomicrographs are sh own in Figure 3 3 and 3 4. The samples have silica contents ranging from 48 to 74 wt. %. Sample LD10 11 is a silica rich (98% silica) igneous rock. As a result, this sample is not considered further for geochemical results. The xenoliths represented by blue diamonds in Figure 3 5 (molar Al/(Na+K) vs. molar Al/(Ca+Na+K) cluster along the metaluminous peraluminous boundary. Trace element geochemistry (Table 3 1 and 3 2) of the meta igneous xenoliths is summarized in Figure 3 6A (contents normalized to p rimitive mantle values of McDonough and Sun,

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78 1995). The plots show enrichments in some fluid mobile incompatible elements, such as Rb, Ba, and Pb, up to 750 times the primitive mantle values. This relative enrichment is paired with minimal enrichment in fluid immobile trace elements, including heavy REE. This is a typical pattern in modern convergent margin rocks (e.g., Thompson et al., 1984). Figure 3 6B shows rare earth element data from the meta igneous xenoliths normalized to the chondritic values o f McDonough and Sun (1995). The samples show variability in the values of Eu and Sr relative to elements of similar compatibility. These include both negative and positive anomalies in Eu and Sr. Nb and Ta are depleted relative to the observed values fo r neighboring elements in all of the samples when normalized to primitive mantle, with the exception of orthogneiss HX 1. U Pb Geochronology O f Zircon Zircon analyses used for age calculations were limited to those with discordance disturbed analyses. For each sample, zircon analyses with ages that do not overlap nsidered discordia which trend to Proterozoic metamorphic events, Cambrian exhumation and exposure, or recent Pb loss. In those samples where there appears to be a meta morphic age, the U concentration in the sample can be examined to discriminate between metamorphic and igneous zircon growth. During metamorphism, fluids can oxidize the U in a rock to the +6 valence state, increasing its solubility (e.g., Rubatto 2002). Because the U is fluid mobile, this leads to any zircons that crystallized under

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79 those conditions to have higher U values. Metamorphic zircons can potentially have higher 238 U cps relative to the mean of the main populations of zircons within a sample. Big Slide diatreme yielded 3 meta igneous samples large enough for whole rock geochemistry and zircon separation (~10 20 cm diameter). BSD10 06 is a garnet granulite and yielded only 2 zircon grains, both of which had discordance grain was 1.74 0.02 Ga and the other was 1.83 0.01 Ga (Table 3 3). BSD10 04 Figure 3 7A). The younger 0% discordant, and with < 3% 206 Pb/ 238 U error, reveals scatter suggestive of disturbance to the zircon U Pb system ( Figure 3 7B). To assist with determining a metamorphic age of the sample, the U intensity in each analysis was examined. Those grains with 238 U cps significantly higher than the mean of the main population (i.e., grain: 1,360,220 cps vs. mean of population: 546,312 cps) may represent metamorphic zircons. Statistically, the mean 207 Pb/ 206 Pb crystallization ages of the high U grains cannot be distinguished from the mean 207 Pb/ 206 Pb crystallization ages of the low U grain, so U intensities are not a discordant grains, including those with high U intensities, yie lds a metamorphic age of 1.72 0.01 Ga. When Hf was analyzed in these zircons, it was observed that there was a very large disconnect between the U Pb ages and the Hf model ages (discussed below), leading to the conclusion that the age of BSD10 04 should be older than the 1.72 Ga mentioned above. Because of this, a single Archean aged grain (2.68 0.4 Ga) with anomalously high error (i.e., 3.9% on 206 Pb/ 238 U) contained in BSD10 04 is believed to be an approximation of the minimum igneous crystallization age for this

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80 sample. Data are summar ized in Table 3 3. BSD10 05 yielded 46 zircons with crystallization age of 1.74 0.01 Ga ( Figure 3 8). Little Sand Creek diatreme yielded 5 meta igneous samples large enough for whole rock geochemistry and zircon separation (~10 20 cm diameter). LSC10 03 and LSC10 10 are amphibolites, LSC10 11 is a mafic gneiss, LSC10 12 is a meta granitoid, and LSC10 13 is a mafic granulite. Data are summarized in Table 3 3. Four of these samples ar e discussed in the following paragraph, but LSC10 12 did not yield a crystallization age and is discussed later. LSC10 03 yielded 49 grains with discordance crystallization age of 1.81 0.01 Ga ( Figure 3 9A). LSC10 10 yielded 26 grains with age of 1.79 0 .1 Ga ( Figure 3 9B). LSC10 13 yielded 48 grains with discordance 206 Pb/ 238 U and 97 Ma on 207 Pb/ 235 U), and 4 outliers were excluded. Forty four grains yielded an upper intercept on concordia of 1 .78 0.03 Ga, which approximates a minimum age for LSC10 13 ( Figure 3 9C). LSC10 Figure 3 10A), including a group of older zircons (~2.75 to 3.14 Ga) which were determined to be xenocrystic based on their Hf v alues, and have been excluded from age calculations. When all of the grains (excluding those interpreted as xenocrysts) are plotted on concordia, the upper intercept of the discordia line is 2.61 0.07 Ga, which is interpreted as an approximate igneous c rystallization age for the sample ( Figure 3 10B).

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81 The lower intercept of the discordia line, interpreted as an approximate metamorphic age is 1.65 0.1 Ga. These data are summarized in Table 3 3. Robinson Ranch diatreme yielded 4 meta igneous samples l arge enough for whole rock geochemistry and zircon separation (~10 20 cm diameter). Data are summarized in Table 3 3. Three of these samples are discussed in the following paragraph, but RRD10 20 did not yield a crystallization age and is discussed later RRD10 05 (meta granitoid) yielded 32 zircons ( age of 1.89 0.01 Ga ( Figure 3 11A). RRD10 09 (meta granitoid) yielded 52 zircons as outliers. Fifty grains yielded a minimum igneous crystallization age of 1.88 0.01 Ga ( Figure 3 11B). RRD10 13 concordia with a single younger grain separated ( Figure 3 12A). The five grains in the older cluster yield an upper intercept age of 1.87 0.01 Ga ( Figure 3 12B), which is interpreted as a minimum igneous crystallization age for the sample. The Lloyd Divide location in the Bearpaw Mountains yielded 1 igneous and 3 meta igneous xenoliths large enough for whole rock geochemistry and zircon separation (~10 20 cm diameter). LD10 01, LD10 07, and LD10 08 are orthogneisses, and LD10 11 is a silica rich igneous rock. Zircons from LD10 07 and LD10 08 revea l two distinct populations of 207 Pb/ 206 Pb ages, which are separated by > 500 Ma and do not overlap at Figure 3 11, 3 12). LD10 07 yielded 55 large measurement error ( i.e., >3% on 206 Pb/ 238 U and 207 Pb/ 235 U). When all of the

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82 grains are plotted on concordia, the upper intercept of the discordia line is 2.62 0.05 Ga, which is interpreted as an approximate igneous crystallization age for the sample ( Figure 3 13). Data a re summarized in Table 3 3. To assist with determining a metamorphic age, 238 U cps were examined. Both of the low discordance analyses contained high 238 U cps relative to the mean of the main population (i.e., grain: 3,455,340 and 6,304,710 cps vs. mean of population: 956,443 cps). A weighted mean of the 2 low discordance analyses yields an age of 1.75 0.19 Ga, which corresponds to documented tectonothermal events in the region and is interpreted as a poorly defined minimum metamorphic age for this sam ple. LD10 08 yielded 60 grains, 32 of which had discordance between 0 and 10% ( Figure 3 14A). This yielded a discordia with an upper intercept of 2.44 0.07 Ga, interpreted to be a minimum igneous crystallization age. When the 30 youngest zircons withi n the sample, (~1.60 to 1.84 Ga) are plotted on concordia, ( Figure 3 14B), it is clear that there is a complex Pb loss pattern within the sample. Seven grains show high 238 U cps relative to the mean of the main population (i.e., grains: 16,126,100 to 10,2 20,000 cps vs. mean of population: 3,030,927 cps). Statistically, the mean 207 Pb/ 206 Pb crystallization ages of the high U cps grains cannot be distinguished from the mean 207 Pb/ 206 Pb crystallization ages of the low U cps grains, so U intensities are not a useful discriminant in this case. When a Ga was calculated. This is interpreted to be a minimum metamorphic age for this sample. LD10 01 yielded 50 grains, only 43 with 15A shows a concordia diagram with all 50 analyses. A discordia calculation ( Figure 3 15A) yielded an upper intercept age of 2.52 0.08 Ga, which is interpreted as a minimum

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83 igneous crystallization age for this sample. Figur e 3 15B shows the cluster of Paleoproterozoic 1.85 Ga. Because of the scatter, 238 U cps were examined. Four grains show high 238 U cps (i.e., grains: 21,978,800 to 7,620,440 cps vs. mean of p opulation: 1,723,033 cps). Statistically, the high U grains cannot be distinguished from the low U grains, so U intensities are not a useful discriminant in this case. A weighted mean of all of the 0.02 Ga, interpreted as a minimum metamorphic age (Table 3 3). LD10 11 yielded 63 grains, 43 of which were used to calculate a minimum igneous crystallization age for the sample of 1.78 0.03 Ga (Table 3 3, Figure 3 16). One orthogneiss xenolith (HX 1) was found within the Highwood Mountains. Data are summarized in Table 3 one of which was an outlier and was exclud ed. In a pattern similar to the orthogneisses from the Grassrange (MX 06, MX 09, MX 10, and MX 11) (Chapter 2), HX 1 yielded an igneous crystallization age as well as an age of metamorphism ( Figure 3 17A). The ielded an upper intercept of 1.85 0.01 Ga, interpreted as a minimum metamorphic age ( Figure 3 17B). The oldest 4 grains similarly yielded an upper intercept of 2.65 0.01 Ga, interpreted as a minimum igneous crystallization age for HX 1 ( Figure 3 17C). The remaining zircon grains yield intermediate ages that lie along a mixing line (discordia) between the older and younger ages.

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84 Samples RRD10 20 (dioritic granulite) and LSC10 12 (meta granitoid) show signs of retrograde metamorphism disturbing the U Pb isotopic system, and complicating any interpretations. Both samples show a common age of zircons from ~1.7 to 1.8 Ga, which can be attributed to possible metamorphism within the GFTZ. The age data are complex and show indeterminate age relationships (Ta ble B 5). Sample RRD10 20 ( Figure 3 18A) yielded a dominant population ranging from ~2.21 to 2.59 Ga, which is likely an igneous population, but Pb loss has affected the sample to the extent that an igneous crystallization age determination would be specu lative at best. There are grains ranging in age from ~2.73 to 3.24 Ga that appear to be from a variety of sources based on their Hf isotopes (discussed below). The older grains appear to lie along a discordia, but do not display Hf systematics indicating they came from the same (or similar) sources, so a discordia regression is not viable. LSC10 12 ( Figure 3 18B) yielded a dominant population of zircons at ~2.8 Ga. There is a scattering of older grains (~3.0 to 3.3 Ga) which appear to be from a variety of sources based on their Hf isotopes (discussed below). LSC10 12 also yielded a population of 14 grains that range in age from 1.92 to 1.68 Ga, which is an indication that the sample was affected by the wide scale metamorphism within the GFTZ. Hf Isoto pe s I n Zircon Zircons were selected for Lu Hf analysis based on different factors. Samples were arranged by 207 Pb/ 206 Pb age to choose a representative population of ages for Hf analysis. Zircons were preferentially chosen for low discordance (as discusse d above) to increase the likelihood of an accurate U Pb age with which to reduce the Hf isotopic data. Zircons were further chosen based on zonation and grain size. If a zircon showed zonation and a U Pb analysis was undertaken in a particular domain, id eally,

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85 the Hf spot would also be within the same domain. The Hf spot size is 40m, and the (initial) calculations used the calculated igneous crystallization ages for the individu al meta igneous samples (Table 3 3). Hf T DM ages are minimum mantle separation ages due to uncertainties in source initial Hf isotopic composition (e.g., Griffin et al., 2002). Two (initial) calculations are discussed: Calculations using t he concordia igneous crystallization age of a given sample, and calculations based on the individual 207 Pb/ 206 (initial) values are recorded in Table 3 3. These values were averaged and a standard deviation a nd median taken to estimate the variance of isotopic values about the mean (Table 3 3). If the median and average are within measurement error of each other, then standard deviation is likely to be a good estimate of the dispersion of the data. The meta igneous samples are plotted on Figure 3 19, ( 207 Pb/ 206 Pb age vs. initial Xenoliths LD10 01, LD10 08, and LSC10 11 yielded zircon ages that followed the pattern seen in orthogneiss xenoliths MX 06, MX 09, MX 10, and MX 11 from the Grassrange (Chapter 2). The samples yielded upper intercept (ig neous crystallization) U Pb ages ranging from 2.44 Ga to 2.61 Ga. Zircon (IA) values for the samples range from 1.8 to 27.7 (Table 3 3), while (2.4 2.6 Ga) values for the samples range from 6.9 to 16.9 (Table 3 (2.4 2.6 Ga) values are likely the result of source compositions ranging from DM like material to evolved (crustal like) material. (2.4 2.6 Ga) values for each sample range from 0.4 to 11.5, with standard deviations of 3.0 to 3.7 (Table 3 3). (i.e., Paleoproterozoic) are similar to those of Archean concordant grains when

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8 6 calculated to the igneous crystallization age of the sample. This suggests that the Paleoproterozoic analyses reflect Pb loss along discordia, and not younger zircon growth, and further suggests that the Hf isotopic system remained primarily closed during metamorphism. These samples give a range of depleted mantle model ages (T DM ) from 2.77 Ga to 3.32 Ga (Table 3 3). Sample LD10 07 a lso yielded ages that followed the pattern seen in orthogneiss xenoliths MX 06, MX 09, MX 10, and MX 11 from the Grassrange (Chapter 2). However, when Hf data were reduced using the upper intercept concordia ages calculated for LD10 07 (~2.62 Ga), certain (2.6 Ga) that plot above DM (observed) values are plotted on a probability density plot ( Figure 3 20A,B), it becomes evident that the Hf data yields a bimodal distribution. This is a clear indication that d uring the ~1.75 Ga metamorphic event new zircon growth integrated Hf with higher (observed) values for these metamorphic (labeled as Figure 3 20A) grains range from 48 to 54 ( Figure 3 20A) as opposed to the magmatic grains (labele Figure 3 20A) that do not appear to have incorporated new Hf which range from 56 to 62 ( Figure 3 20A). Similarly, the Hf depleted mantle model ages (Hf T DM ) also appear to be bimodal ( Figure 3 20B) with averages of 2.62 and 2.88 Ga and st andard deviations of 0.06 and 0.04 for the metamorphic and magmatic respectively for LD10 07. BSD10 04 primarily yielded Paleoproterozoic age zircons. When (1.8) was (1.8) = ~ 30; Table 3 3), and Hf T (DM) ages were calculated and yielded ages of ~3.4 Ga, indicating a much older crustal (2.7) was calculated

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87 for the sample based on the single U Pb analysis mentioned above, the calculated values shifted from ~ 30 to 12 which fits the evolution model of the Archean age xenoliths better than the values calculated at ~1.8 Ga. When plotted on Fig ure 3 19, the sample lies along a reasonable Lu Hf crustal evolution curve indicating the sample was primarily a crustal melt. (2.7 Ga) ranged from 9.4 to 14.0 (Table 3 3), (IA) values ranged from 30.9 to 37.7 (Table 3 3), and BSD10 04 average zircon Hf T (DM) calculated to 3362 Ma. For the Paleoproterozoic samples (LSC10 03, LSC10 10, LSC10 13, RRD10 05, RRD10 0 9, RRD10 13 and LD10 11) the measured Hf isotopic compositions for all zircons were recalculated to the U (T) ), which can be seen in Table 3 (1.8 1.9 Ga) the Paleoproterozoic samples range from 13.8 to 8.0 with averages between 9.4 to 6.3 and standard deviations between 0.8 to 3.0 (Table 3 3). BSD10 06 only yielded 2 grains eligible for Hf analysis, and each grain was reduced to its individual 207 Pb/ 206 Pb a ge (IA) (IA) of 10.1 and 7.0. Average Hf T (DM) model ages for the Paleoproterozoic samples (BSD10 06, LD10 11, LSC10 03, LSC10 10, LSC10 13, RRD10 05, and RRD10 09, RRD10 13) range from ~1.84 Ga in BSD10 06 to ~2.59 Ga in RRD10 13 (Table 3 3). The Hf T DM ages represent minimum mantle separation ages because Lu/Hf in zircon is invariably lower than in whole rocks (e.g., Griffin et al., 2002). Calculation of secondary or crustal residence ages requires knowledge of the Lu/Hf of the source(s) and were not calculated because the data ultimately suggest a mixing of crustal and mantle sources at ~1.8 to 1.9 Ga.

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88 RRD10 20 yielded an older population of zircons (~2.7 to 3.2 Ga) which appear to lie along a discordia which yields an upper intercept age of ~3.4 Ga. However, when (3.4) the isotope values for the older population of zircons lie above values for DM. This indicates that these older grains likely do not come from the same source and should not be regressed as a g roup, and that the oldest component within RRD10 20 is not ~3.4 Ga. The ~2.7 to 3.2 Ga grains may have been acquired from multiple sedimentary sources. Similar to RRD10 20, LSC10 12 yielded an older population of zircons (~3.0 to 3.3 Ga) which appear to lie along a discordia that yields an upper intercept age of ~3.4 Ga. However, when ~3.4 Ga is used for calculating (3.4) for LSC10 12, the isotope values for the older population of zircons lie above values for DM. This indicates that these older grains likely do not come from the same source and should not be regressed as a group, and that the oldest component wit hin LSC10 12 is not ~3.4 Ga. Sm Nd Whole Rock Isotopes As a group, the meta (0) from 5.6 to 45.9. Initial ratios were calculated using the best estimates of the individual rock crystallization ages ( 207 Pb/ 206 Pb age s 1.79 Ga to 2.68 Ga) shown in Table 3 4. This (2.4 2.7) of 0.3 to 5.0 (Table 3 4, Figure 3 (1.7 1.9) of 0.9 to 10.0 (Table 3 4, Figure 3 21). Depleted mantle model ages for the Archean meta igneous samples were calculated usin g the model of DePaolo (1981) and ranged from 2.78 Ga to 3.33 Ga ( Figure 3 22) (excluding LSC10 11). The Paleoproterozoic meta igneous samples (excluding LD10 11) yielded depleted mantle model ages (DePaolo, 1981) between 2.46 Ga and 2.55 Ga (Table 3 4), overlapping with the igneous crystallization age of the ~2.4 to 2.6 Ga meta igneous samples. RRD10 13 yielded an igneous

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89 crystallization age of 1.87 Ga, and a Nd T (DM) of 3.06 Ga, which is far older than the other Paleoproterozoic samples. This indicates that the source melt for RRD10 13 was values of all of the meta igneous samples (excluding LD10 11) are shown ( Figure 3 21) es (green bar) from the Little Belt Mountains northern Wyoming province (Miller et al., 1986; Wooden and Mueller, 1988; Frost, 1993; Mueller et al., 1996; Mueller et a l., 2004; Frost et al., 2006; Mueller et al., 2010). Shown on Figure 3 21, two evolution lines approximating lower continental crustal evolution using 147 Sm/ 144 Nd = 0.1532 (Rudnick and Gao, 2003) are plotted. These lines begin at DM and CHUR/BSE values a t 3.5 Ga, based on the oldest Sm Nd T DM ages derived from the xenoliths Isotopically, the Archean meta (2.4 2.7 Ga) values range 147 Sm/ 144 Nd = 0.1532; (2.5 Ga) = 10 2), suggesting differing levels of involvement of b oth crustal and mantle sources ( Figure 3 21). The Proterozoic meta (1.7 1.9 Ga) values that also range between DM and lower crustal values, indicating different degrees of mixing between juvenile and evolved sources ( Figure 3 21 ). Sample LSC10 11 yielded a Sm Nd T (DM) of 4.96, and LD10 11 yielded a Sm Nd T (DM) of ~1.59 Ga, which is younger than the calculated igneous crystallization age of ~1.83 Ga for the sample. These values are not geologically realistic, and are most likely attributable to difficulties in measuring such low concentrations of Sm and Nd.

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90 Discussion Origins O f T he Meta I gneous Xenoliths The xenoliths from the Missouri Breaks Diatremes, the Bearpaw Mountains, and the Highwood Mountains show similarities in trac e elements, U Pb ages, and Hf and Sm Nd isotopic systems to those from further south in the GFTZ in the Grassrange area (Chapter 2). Zircon U/Pb ages from plutonic and meta igneous xenoliths yield two suites of igneous crystallization ages: an older range between 2.43 Ga and 2.68 Ga (six samples), and younger ages between 1.74 Ga and 1.89 Ga (eight samples). When Hf isotopic data within each sample are reduced to the sample igneous crystallization age, each suite of zircons shows a relatively small spread error envelopes (excluding LD10 07). This suggests that the younger zircons within the samples experienced a Pb loss event(s) that effected their U Pb ages, but did not affect their Hf isotopes. This indicates that the gra ins affected by Pb loss formed from magmas with the same Hf isotopes as the concordant grains. The range of initial Hf isotopic compositions for the Paleoproterozoic samples scatter from DM to slightly below CHUR/BSE, which are values typical of an arc ty pe environment (e.g., Dhuime et al., 2011). Trace element patterns are also suggestive of a subduction zone origin or recycling of subduction generated crust. This includes the relatively elevated LIL and suppressed HFSE contents commonly observed in arc related rocks. Enrichment in fluid mobile elements (e.g., Pb and Ba) relative to immobile elements is a common effect of fluid fluxing in subduction environments (Pearce, 1983; Thompson et al., 1984). Tectonic discrimination diagrams ( Figure 3 23; Pearce 1983) also support a volcanic arc origin for the protoliths of the samples. Some of the older meta igneous samples have positive Eu and Sr anomalies indicating accumulation of plagioclase, but others

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91 show negative Eu and Sr anomalies ( Figure 3 6B), which indicates the removal of plagioclase during petrogenesis by fractional crystallization or retention of plagioclase in the source. Very minor depletion in the HREE is evident in some samples (Fig 3 6A), which is likely due to residual garnet in the source These data suggest that the Paleoproterozoic meta igneous xenoliths represent magmas formed and crystallized during the ocean subduction and subsequent continent continent collision that formed the GFTZ. Trace element patterns for the Paleoproterozoic samples are very similar to the pattern observed in the older meta igneous samples ( Figure 3 6), suggesting both may have originated in a subduction zone environment, or recycled such crust. Whole rock Nd and zircon Hf isotopic data reveal different mixe s of mantle and crustal material. Figure 3 from all xenoliths using the apparent 207 Pb/ 206 Pb age of each individual zircon for the (2.4 2.6 Ga) (1.7 1.9 Ga) for the meta igneou s samples. The older samples can be split into two groups. The first, LD10 01, LD10 07, and LD10 08, crystallized between ~2.44 Ga and ~2.62 Ga, but have mean Hf T DM for their zircons from 2.77 Ga to 2.88 Ga and (2.4 2.6 Ga) values from 9.8 to 6.9 ( F igure 3 19) The (2.4 2.6 Ga) values for the samples are suggestive of depleted mantle playing an important role in their petrogenesis with ~2.4 to 2.6 Ga ( Figure 3 19) However, (2.4 2.6 Ga) values lie below CHUR/BSE which suggests mixing between depleted mantle and an old enriched crustal component occurred in the late Archean and/or early Paleoproterozoic ( Figure 3 21 Mueller et al., 2010; Mirnejad and Bell, 2006; Mirnejad and Bell, 2008). The evolve d source may have been involved via direct assimilation of c. 2.8 Ga or older Wyoming craton materials, or

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92 that subducted sediments contributed this evolved component (Mueller et al., 2010). The second group includes Archean xenoliths BSD10 04 and LSC10 1 1, which exhibit much lower (2.6 2.7 Ga) ranging from 5.7 to 16.9. This indicates that the samples crystallized from a primarily evolved (crustal) melt with only minor input from a more juvenile source (DM). While LSC10 BSD10 04 is 5.0 which lies below CHUR/BSE and is another indication of a greater input of evolved (crustal) material, such as older MHB crust. The Paleoproterozoic samples (~1.74 Ga to ~1.89 Ga, BSD10 05, LSC10 03, LSC10 10, LSC10 13, RRD10 05, RRD10 09, RRD10 13, and LD10 11) have average Hf T DM ages of their zircons of ~1.95 Ga to ~2.59 Ga (Table 3 3), which suggests differing influences of older crust and depleted mantle in their genesis. Shown on Figure 3 19, two evolution line s approximating lower continental crustal evolution using 176 Lu/ 177 Hf = 0.0187 (Rudnick and Gao, 2003) are plotted. These lines begin at DM and CHUR/BSE values at 3.5 Ga, based on the oldest Hf T DM ages derived from the xenoliths (1.7 1.9 Ga) values for the Paleoproterozoic samples cluster well above the evolution line. (1.7 1.9 Ga) range close to DM values at ~1.7 to 1.9 Ga, which suggests that depleted mantle played an important part of their pe trogenesis ( Figure 3 19) 9.6) are similar to those of similar age from the Little Belt Mountains (Weiss et al., 2009) ( Figure 3 19). While some of the (1.8 1.9 Ga) values are more negative than those from the Little Belt Mou ntains, both data sets seem to indicate a mixture between a DM like source and an older crustal component to different degrees, likely formed in an arc like environment. The single exception to this pattern is BSD10 05 which has an igneous crystallization age of 1.74

PAGE 93

93 Ga, similar to the ages of the granitoid samples from the Grassrange (Chapter 2). The range of (1.7) values for the sample is from 6.3 to 11.0, which overlaps with the values from the Grassrange, indicating that BSD10 05 has a larger influence of crustal 0.9 t o 7.4 ( Figure 3 03 and LSC10 and reinforce the suggestion that the samples might have formed in an ocean continent subd 05, LSC10 13, RRD10 05, RRD10 09, and RRD10 13) are very similar to the ~1.7 Ga granitoid samples (MX 08, MX 18) from the Grassrange (Chapter 2) and seem to indicate slightly more mixing of an older crustal component within these samples, possibly indicating primary formation after oceanic subduction was completed and during the ensuing continental collision. Implications F or Great Falls Tectonic Zone Evolution The dataset presented in th is study is best interpreted in conjunction with the existing data from the Grassrange area to the south (Chapter 2) and the Little Belt Mountains to the southwest (Mueller et al., 2002). When taken together, these studies provide an enhanced characteriza tion of the poorly exposed GFTZ crust in central Montana. Evidence from seven of the eight meta igneous Paleoproterozoic samples (LSC10 03, LSC10 10, LSC10 13, RRD10 05, RRD10 09, RRD10 13, and LD10 11) suggest that these rocks formed 1.79 1.89 Ga ago, an d that similar to the Little Belt Mountains (Mueller et al., 2002; Weiss et al., 2009) the depleted mantle was an important part of their petrogenesis. The relatively juvenile initial Hf and Nd compositions ( Figure 3 19; Figure 3 21) indicate that the sa mples were most likely

PAGE 94

94 formed during the consumption of oceanic crust between the Medicine Hat block and the northern Wyoming craton. These data support models proposed by Mueller et al. (2002, 2005) that the GFTZ resulted from ocean basin closure between the Wyoming Craton and Medicine Hat block (see also Harms et al., 2004). The primarily (1.7) (1.7) (whole rock) values for Paleoproterozoic sample BSD10 05 ( Figure 3 19; Figure 3 21) both indicate substantial mixing betw een a DM like component and an evolved Archean crustal component to a greater extent than evident in the slightly older (~1.79 1.89 Ga) Missouri Breaks xenoliths and igneous rocks in the Little Belt Mountains. This suggests that the ~1.74 Ga BSD10 05 in teracted with a significant amount of evolved material during its petrogenesis, a process consistent with formation in a continental collision zone in which older crust was thickened and melted. The trace element patterns ( Figure 3 6) from the Archean and Paleoproterozoic orthogneisses (i.e., relative depletion of HFSE) are also consistent with patterns in rocks from modern continental arcs (e .g., Thompson et al., 1984; Pearce et al., 1984 ); however, this does not discriminate between volcanic arc, syn col lis i onal, or post collisional origins. Isotopic data from the Little Belt Mountains (Mueller et al., 2002) and from this study indicate that the igneous arc was built upon continental crust. (1.7 1.9 Ga) (1.7 1.9 Ga) (whole rock) values for the Paleoproterozoic meta igneous samples ( Figure 3 19; Figure 3 21) suggest that mixing between a DM like component and an evolved Archean crustal component occurred to a slightly great er extent than is observed in the (~1.86 Ga) Little Belt Mountain sample suite (Mueller et al., 2002; Weiss et al., 2009).

PAGE 95

95 Geochronologic data provide insight into the cratonic affinity of the igneous arc basement. Figure 3 24, illustrates the ranges of ages observed in this study compared to published data from the Wyoming craton (Heimlich and Banks, 1968; Mueller et al., 1995; Kirkwood, 2000; Mogk et al., 2004; Frost and Fanning, 2006; Mueller et al., 2010, Krogh et al., 2011) and Medicine Hat block (V illeneuve et al., 1993; Davis et al., 1995; Gorman et al., 2002). Ages were limited to those older than 2.4 Ga in order to make the most direct comparisons of pre GFTZ data. Igneous crystallization ages of the meta igneous samples (indicated by a star) a re plotted in Figure 3 24. The dominance of ~2.4 2.7 Ga igneous crystallization ages suggests that the meta igneous samples share a closer affinity to the MHB, which is characterized by Neoarchean ages, than to the northern Wyoming Province, which is do minantly Mesoarchean ( Figure 3 24). Based on the affinity of the older meta igneous samples from this study the Paleoproterozoic magmatic arc was built on MHB crust. This implies the polarity of ocean subduction was beneath the MHB, which is supported by the Deep Probe velocity model of Gorman et al. (2002). These authors cited an upper mantle floating reflector, interpreted to be a subducted slab, dipping beneath the MHB. During subduction, the Wyoming craton would be attached to the down going slab, l ikely leading to MHB crust over riding Wyoming crust. Further evidence of this can be seen in SW Montana within the Tobacco Root Mountains, which preserves evidence of Wyoming province material being buried to depths great enough for anatectic melts to ha ve been created at 1.77 Ga (Mueller et al., 2004, 2005). Harms et al. (2004) calculated pressure temperature time paths from three metamorphic suites and metamorphosed mafic dikes and sills from the Tobacco Root

PAGE 96

96 Mountains which revealed crustal thickenin g, metamorphism, and partial melting of rock occurring between 1.78 and 1.76 Ga. Similarly, Mueller et al. (2005) showed evidence from zircons from the Tobacco Root Mountains and monazite from the Highland Mountains that the northwestern Wyoming province experienced an episode of high grade metamorphism at ~1.77 Ga. Further, Roberts et al. (2002) examined rocks from the Montana meta sedimentary terrane (MMT, which includes the Tobacco Root Mountains and the Highlands), which yielded 430 40 Ar/ 39 Ar biotite ages as well as 4 207 Pb/ 206 Pb step leached garnet ages. Roberts et al. (2002) suggested that metamorphic events in the MMT took place between 1.82 and 1.79 (garnet growth ages), and that between 1.78 to 1.74 Ga the MMT was rapidly cooling (biotite cooling ages). We suggest that the 1.78 to 1.77 Ga metamorphic ages are representative of terminal collision between the MHB and Wyoming province, and that this lead to thickening and crustally derived granitoid magmatism represented by the ~1.73 and 1.74 Ga gra nitoids from the Grassrange diatremes (Chapter 2), as well as metamorphism of the MHB country rock represented by the 1.75 Ga metamorphic ages of the orthogneiss samples from the Bearpaw Mountains (this study). The timing of arc magmatism and continental c ollision in the GFTZ broadly coincides with the overall amalgamation of the constituent Laurentian cratons (Mueller et al., 2011; Figure 3 25). These include collisions between the Wyoming Superior and Hearn Superior, which formed the Trans Hudson orogen at 1.7 1.9 Ga ( Karlstrom and Houston, 1984; Lewry et al., 1987; Hoffman, 1988; Bickford et al., 1990; Gordon et al., 1990; Machado, 1990; Nelson et al., 1993; Lewry et al., 1994; Lucas et al., 1996; Resor et al., 1996; Sun et al., 1996; Dahl et al., 1999 ; Machado et al., 1999; Mueller et al.,

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97 2002 ), and the Wyoming Colorado collision at ~1.66 1.80 Ga ( Bickford and Boardman, 1984; Sims and Peterman, 1986; Karlstrom and Bowring, 1988; Premo and Van Schmus, 1989; Van Schmus et al., 1993; Chamberlain, 1998; Selverstone et al., 2000; Hill and Bickford, 2001; Hill, 2004 ). Data from this study reinforce the generally synchronous nature of the amalgamation event and places significant limits on plate geometries during Laurentian growth (e.g., Hoffman, 1989; Mue ller et al., 2005). Conclusions Zircons in the meta igneous samples fall into two categories. The first category shows two distinct age populations. The older population is represented by magmatic zircons with upper intercept concordia ages of ~2.43 Ga to ~2.65 Ga. These same samples contain a younger population of metamorphic zircons with minimum ages of ~1.72 Ga to ~1.85 Ga. The Archean igneous crystallization ages documented in the xenoliths is consistent with younger ages observed in the MHB. These ages are younger than those expected from the northern Wyoming Craton, indicating that the Paleoproterozoic arc was built upon MHB Archean crust. The second category of meta igneous samples shows a single age population represented by magmatic zirco ns with upper intercept concordia ages of ~1.74 Ga to ~1.89 Ga. Geochemistry and geochronology of the metamorphic ages of zircons in the Archean samples and the igneous crystallization ages of the Paleoproterozoic samples correlate with the ~1.86 Ga magma tism in the Little Belt Mountains and the tectonic events associated with the collision of the Medicine Hat block with the Wyoming craton that terminated formation the Great Falls tectonic zone. The geochemistry of the crustal xenoliths from this study su pport the proposal that the GFTZ experienced ocean subduction and arc magmatism from 1.89 to 1.78 Ga. These data also suggest that the volcanic arc created during

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98 ocean continent subduction was built on MHB crust, rather than that of the Wyoming Province, and that the polarity of subduction was beneath the MHB. Evidence from the southwestern GFTZ suggests that the system transitioned to continent continent collision around 1.78 Ga to 1.77 Ga when there was a period of crustal thickening and metamorphism w ithin the GFTZ related to terminal collision. After terminal collision was completed, there was then post orogenic collapse and extension which likely resulted in upwelling mantle heating and partially melting the lower crust leading to crustally derived granitoid magmatism This can be seen in the ~1.73 and 1.74 Ga granitoid xenoliths from the Grassrange (Chapter 2) as well as ~1.74 Ga mafic meta granitoid sample BSD10 05 from the Missouri Breaks. The G reat F alls tectonic zone therefore, is constrained to the same general 150 Ma time period during which other blocks were sutured to form Paleoproterozoic Laurentia (Hoffman, 1988, 1989).

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99 Table 3 1. Major, trace and rare earth element dat a of igneous protolith xenolith s from the B ig Slide and Little Sa nd Creek d iatremes. Major elements in wt. %, trace and rare earth elements in ppm. Sample BSD10 04* BSD10 05* BSD10 06 LSC10 03 LSC10 10* LSC10 11* LSC10 12* LSC10 13* wt. % SiO 2 73 48 60 50 48 57 69 48 TiO 2 0.09 4.7 0.60 0.55 0.9 0 0.16 BDL 0.57 Al 2 O 3 15 14 16 18 21 15 16 8.0 Fe 2 O 3 0.69 18 8.1 7.8 6.1 1.1 0.47 5.2 MnO 0.04 0.05 0.15 0.13 0.09 0.02 0.02 0.18 MgO 0.21 12.8 3.7 6.2 3.1 2.2 0.08 14 CaO 2.2 1.8 4.6 6.9 13 0.96 4.4 14 Na 2 O 5.7 1.5 3.9 3.9 1.7 0.54 2.3 1.4 K 2 O 2.9 6.8 0.81 2.9 0.75 10 9.7 2.2 P 2 O 5 0.04 0.39 0.12 0.27 0.12 0.01 0.06 0.32 LOI 0.88 0.93 1.4 3.1 1.5 2.0 2.2 1.7 Total 101.31 108.87 99.71 99.76 96.07 88.45 104.18 95.66 ppm Li N/A 43 N/A 7.8 4.5 N/A N/A 13 Sc N/A 4.8 N/A 26 15 N/A N/A 39 Ti N/A 23187 N/A 3418 5922 N/A N/A 3814 V N/A 284 N/A 114 192 N/A N/A 115 Cr N/A 50 N/A 209 49 N/A N/A 16 Co N/A 37 N/A 32 20 N/A N/A 12 Ni N/A 85 N/A 120 37 N/A N/A 11 Cu N/A 4.4 N/A 16 45 N/A N/A 11 Zn N/A 249 N/A 85 38 N/ A N/A 283 Ga N/A 39 N/A 16 30 N/A N/A 14 Rb N/A 323 N/A 103 34 N/A N/A 110 Sr N/A 278 N/A 1644 1547 N/A N/A 410 Y N/A 2.2 N/A 37 12 N/A N/A 28 Zr N/A 38 N/A 25 31 N/A N/A 49 Nb N/A 9.8 N/A 5.0 7.0 N/A N/A 11 Cs N/A 1.6 N/A 4.3 0.31 N/A N/A 2.8 Ba N/A 6863 N/A 5724 829 N/A N/A 1365

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100 Table 3 1. Continued. Sample BSD10 04* BSD10 05* BSD10 06 LSC10 03 LSC10 10* LSC10 11* LSC10 12* LSC10 13* ppm La N/A 28 N/A 21 7.3 N/A N/A 11 Ce N/A 55 N/A 55 17 N/A N/A 32 P r N/A 6.2 N/A 7.9 2.2 N/A N/A 5.1 Nd N/A 23 N/A 35 9.0 N/A N/A 26 Sm N/A 3.0 N/A 8.1 2.4 N/A N/A 8.7 Eu N/A 1.4 N/A 1.9 1.1 N/A N/A 1.3 Gd N/A 2.0 N/A 7.2 2.3 N/A N/A 8.1 Tb N/A 0.16 N/A 1.0 0.35 N/A N/A 1.2 Dy N/A 0.49 N/A 6.2 2.1 N/A N /A 6.0 Ho N/A 0.06 N/A 1.2 0.41 N/A N/A 0.96 Er N/A 0.19 N/A 3.3 1.0 N/A N/A 2.3 Tm N/A 0.01 N/A 0.49 0.18 N/A N/A 0.30 Yb N/A 0.06 N/A 3.0 1.1 N/A N/A 1.6 Lu N/A 0.01 N/A 0.40 0.16 N/A N/A 0.20 Hf N/A 0.30 N/A 0.46 1.1 N/A N/A 1.3 Ta N /A 0.23 N/A 0.16 0.47 N/A N/A 0.94 Pb N/A 7.0 N/A 11 3.4 N/A N/A 9.7 Th N/A 0.28 N/A 0.37 1.9 N/A N/A 6.6 U N/A 0.17 N/A 0.90 0.49 N/A N/A 1.7 BDL below detection limit N/A not applicable *Major elements were analyzed with an incorrect c alibration, thus values may be off by 10%.

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101 Table 3 2. Major, trace and rare earth element data of xenoliths from Robinson Ranch Diatreme, the Bearpaw Mountains, and the Highwood Mountains. Major elements in wt. %, trace and rare earth elements in ppm. Sample RRD10 05* RRD10 09* RRD10 13* RRD10 20* LD10 01 LD10 07* LD10 08 LD10 11 HX 1 wt. % SiO 2 58 58 56 64 74 70 68 98 67 TiO 2 0.65 0.65 0.80 0.01 0.15 0.32 0.42 0.05 0.69 Al 2 O 3 11 13 20 22 14 16 14 0.77 14 Fe 2 O 3 6.2 6.1 9.2 0.49 2.5 6.6 5.2 1.1 7.2 MnO 0.09 0.09 0.08 0.02 0.04 0.08 0.06 0.01 0.10 MgO 8.5 8.6 6.3 1.2 0.81 1.9 1.9 0.03 1.6 CaO 2.8 4.9 0.97 2.1 2.1 2.5 2.1 0.06 4.5 Na 2 O 1.9 2.2 1.2 4.4 3.5 3.8 3.0 0.40 3.0 K 2 O 3.5 3.6 7.8 7.9 2.6 3.6 3.3 0.06 4.1 P 2 O 5 0.19 0.27 0.03 0.04 0.02 0.06 0.02 0.03 0.30 LOI 0.46 0.59 0.59 0.51 0.23 0.37 1.3 0.27 0.23 Total 93.67 98.48 103.59 103.00 99.93 105.17 99.67 100.38 102.76 ppm Li 29 24 7.2 4.9 4.7 5.7 4.7 1.5 5.9 Sc 13 13 23 0.29 5.2 11 11 0.27 12 Ti 3220 3346 5284 661 872 2665 2717 452 3644 V 91 98 215 5.5 32 73 74 8.7 34 Cr 158 217 285 3.7 35 94 102 6.3 25 Co 19 20 20 0.83 5.1 14 11 0.57 6.0 Ni 81 92 46 5.1 13 42 14 3.6 36 Cu 18 6.4 37 1.1 10 32 7.6 24 22 Zn 61 60 83 4.6 26 6 0 40 4.9 141 Ga 17 17 24 17 15 18 17 BDL 23 Rb 106 95 308 198 42 60 59 0.67 89 Sr 685 781 366 1554 447 304 281 63 386 Y 14 14 26 1.5 9.5 21 22 0.68 40 Zr 10 35 56 16 5.4 16 35 1.4 100 Nb 5.2 4.3 10 BDL 0.63 4.0 6.8 BDL 17 Cs 1.4 0.82 3 .3 0.19 0.46 0.39 0.75 0.14 0.34 Ba 3507 5239 5086 15074 1577 1709 1363 BDL 3462

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102 Table 3 2. Continued. Sample RRD10 05* RRD10 09* RRD10 13* RRD10 20* LD10 01 LD10 07* LD10 08 LD10 11 HX 1 ppm La 47 40 35 6.9 23 49 18 0.74 53 Ce 90 77 59 12 36 83 28 1.4 110 Pr 9.8 9.1 6.2 1.2 3.3 8.3 2.7 0.18 12 Nd 38 35 22 4.0 9.1 30 9.3 0.72 49 Sm 6.1 6.0 4.4 0.67 1.3 4.8 2.1 0.13 8.9 Eu 1.9 2.0 2.2 2.3 1.1 1.4 0.89 0.04 2.7 Gd 4.4 4.7 4.3 0.81 1.5 4.4 2.8 0.12 8.1 Tb 0.5 3 0.53 0.68 0.05 0.24 0.65 0.56 0.01 1.2 Dy 2.6 2.5 4.1 0.21 1.6 3.7 3.5 0.09 6.8 Ho 0.47 0.45 0.81 0.03 0.34 0.72 0.68 0.02 1.3 Er 1.2 1.3 2.4 0.09 0.95 2.0 1.8 0.05 3.8 Tm 0.20 0.17 0.36 0.01 0.19 0.32 0.27 0.01 0.52 Yb 1.1 1.0 2.3 0.06 1.1 2.0 1.6 0.04 3.3 Lu 0.17 0.15 0.34 0.01 0.17 0.30 0.23 0.01 0.48 Hf 0.49 0.61 1.2 BDL 0.21 0.51 0.77 BDL 2.3 Ta 0.13 0.11 0.67 0.03 0.02 0.05 0.10 BDL 0.79 Pb 19 18 28 11 26 28 29 3.81 23 Th 1.3 0.05 7.2 1.4 0.42 6.5 1.1 BDL 5.7 U 0.15 0 .18 1.3 0.45 0.08 0.13 0.33 0.22 0.30 BDL below detection limit. *Major elements were analyzed with an incorrect calibration, thus values may be off by 10%.

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103 Table 3 3. Igneous p rotolith xenoliths, Missouri Breaks, zircon LA ICP MS U Pb data 207 Pb / 206 Pb ages and Hf isotope data reported (Ma). Sample 207 Pb/ 206 Pb Error 207 Pb/ 206 Pb Error (IA) (IA) (T) (T) (T) (T) (T) Hf Std. Number Age Age Highest Lowest Highest Lowest Mean S.D.** Median T (DM) a Dev. Big Slide Diatreme BSD10 04 2684 39 1718 12 30.9 37.7 9.4 14.0 11.3 1.2 11.2 3362 46 B SD10 05 N/A N/A 1738 12 6.5 11.7 6.3 11.0 9.2 1.2 9.2 2466 45 BSD10 06 1742 24 1831 10 10.1 7.0 N/A N/A N/A N/A N/A 1844 19 Little Sand Creek Diatreme LSC10 03 N/A N/A 1805 4 8.0 4.1 8.0 4.1 6.3 0.8 6.4 1945 31 LSC10 10 N/A N/A 1790 1 3 2.8 3.9 2.3 3.9 1.5 1.6 1.5 2237 64 LSC10 11 2606 65 1647 100 2.6 24.1 5.7 16.9 11.5 3.0 11.1 3319 119 LSC10 13 N/A N/A 1783 25 3.7 9.5 4.4 9.1 6.2 1.2 6.1 2396 43 Robinson Ranch Diatreme RRD10 05 N/A N/A 1894 11 3.2 8.9 3.1 8.7 6.1 1.3 6.2 2492 49 RRD10 09 N/A N/A 1875 8 3.1 9.9 3.5 9.6 5.6 1.4 5.4 2453 53 RRD10 13 N/A N/A 1872 10 3.3 19.0 3.4 13.8 9.4 3.0 9.5 2589 110 Lloyd Divide, Bearpaw Mountains LD10 01 2517 76 1738 22 7.5 24.3 6.9 4.1 0.4 3.5 1.3 2810 132 LD10 07 2622 54 1746 190 4.2 19.1 1.2 1.6 0.3 1.0 0.8 2878 36 ***Metamorphic zircons 10.0 15.43 13.0 1.7 13.1 2616 64 LD10 08 2438 70 1718 28 1.8 24.7 4.6 9.8 1.1 3.7 1.6 2773 141 LD10 11 N/A N/A 1777 33 6.3 4.8 6.2 2.2 2.0 2.5 2.1 2085 95 Highwood Mountains HX 1 2649 8 1846 50 N/A N/A N/A N/A N/A N/A N/A N/A N/A (IA) zircons reduced to individual U Pb age. *zircons reduced to 207 Pb / 206 Pb age, excluding inherited grains. a DM model ages were c alculated using the model of Mueller et al. (2008). **Standard deviation. ***Metamorphic zircons from sample LD10 07, reduced to 1.75 Ga.

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104 Table 3 4. LA ICP MS Nd isotope data reported (Ma). Sample Age (Ma) Sm Nd 143 Nd/ 144 Nd (0) (T) Nd T (DM) a Number (ppm) (ppm) Big Slide Diatreme BSD10 04 2684 1.1 8.6 0.510276 45.9 5.0 3129 BSD10 05 1738 3.0 22.7 0.510896 33.8 8.2 2470 Little Sand Creek Diatreme LSC10 03 1803 8.1 34.8 0.511883 14.6 1.9 2459 LSC10 10 1790 2.4 9.0 0.512158 9.2 0.9 2540 LSC10 11 2606 0.3 1.6 0.510508 41.4 N/A N/A LSC10 12 N/A 0.6 2.3 0.511122 29.4 N/A N/A LSC10 13 1783 8.7 25.7 0.512343 5.6 7.4 N/A Robinson Ranch Diatreme RRD10 05 1890 6.1 38.0 0 .511126 29.3 5.3 2508 RRD10 09 1875 6.0 35.3 0.511188 28.1 5.6 2550 RRD10 13 1872 4.4 22.3 0.511176 28.4 10.0 3059 RRD10 20 N/A 0.7 4.0 0.510820 35.3 N/A 3075 Lloyd Divide, Bearpaw Mountains LD10 01 2517 1.3 9.1 0.510762 36.4 1.4 27 83 LD10 07 2430 4.8 29.6 0.510920 33.4 3.1 2848 LD10 08 2551 2.1 9.3 0.511319 25.6 4.9 3332 LD10 11 1832 0.1 0.7 0.511955 13.2 N/A 1589 Highwood Mountains HX 1 2649 7.6 40.2 0.511106 33.2 0.3 3265 *Orthogneiss and amphibolite sampl es reduced to 207 Pb / 206 Pb age (Table 3 3 ). a DM model ages were calculated using the model of DePaolo (1981).

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105 Figure 3 1. Generalized map of Precambrian basement provinces of southwestern Laurentia (after Ross et al., 1991; Condie, 1992; Doughty et a l., 1998; Vogl et al., 2004; Foster et al., 2006 2012 ). Exposures of basement in Laramide style uplifts are shown in the dark grey shaded areas.

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106 Figure 3 2. Generalized depiction of Cenozoic alkaline rock occurrences in the Montana alkali province ( after Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well defined and are based on aeromagnetic data from Sims et al. (2004). Sample locations shown as red stars.

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107 Figure 3 3. Photomicrographs of selected met a igneous samples, with Plane Polarized light (PPL) images on the left, and Cross Polarized Light (XPL) images on the right. Scale bar is 1 mm. (A) Sample LD10 01, a garnet quartzofeldspathic gneiss. Compositional banding is defined by garnet rich and q uartzofeldspathic layers. Note that both garnet and plagioclase are sub to euhedral and fresh. (B) Sample RRD10 05, an amphibole biotite meta granitoid. Amphibole occurs as, anhedral, altered grains with biotite reacting with the amphibole. (C) Sample LSC10 03 is a fine to medium grained lineated biotite amphibolite, where amphibole grain preferred orientation defines the rock fabric.

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108 Figure 3 4. Photomicrographs of selected meta igneous samples, with Plane Polarized light (PPL) images on the left and Cross Polarized Light (XPL) images on the right. Scale bar is 1 mm. (A) BSD10 05 is a mafic meta granitoid. Biotite is the dominant phase, with plagioclase, orthoclase and quartz common as well. (B) LSC10 11 and (C) LSC10 12 are presumed highly a ltered and retrogressed meta igneous samples. See text for details.

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109 Figure 3 5. Xenolith samples plotted on an alumina saturation index (Shand, 1943; modified by Frost et al., 2001). Meta igneous blue diamonds.

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110 Figure 3 6. (A) Primitive mantl 1995) for meta igneous samples. (B) Chondrite normalized REE diagram (McDonough and Sun, 1995) for meta igneous samples using data in Table 3 1 and 3 2. Relative enrichment of LILEs to HFSE suggests form ation in an arc environment.

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111 Figure 3 7. (A) Concordia diagram plotting composite U Pb data for meta granitoid sample BSD10 error. (B) Expanded view of lower end of concordia showing sprea d of samples and weighted mean age.

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112 Figure 3 8. Concordia diagram plotting composite U Pb data for mafic meta granitoid sample BSD10 error.

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113 Figure 3 9. Concordia diagram showing U Pb up per intercept regression for amphibolite samples LSC10 03 (A) and LSC10 10 (B) as well as mafic granulitic gneiss LSC10 13 (C). Each ellipse represents a single spot

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114 Figure 3 10. (A) Concordia diagram plotting discordance Pb data for mafic gneiss sample LSC10 11. Each ellipse represents a single spot analysis and showing spread of samples and the calculated upper intercept age.

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115 Figure 3 11. Concordia diagram showing U Pb upper intercept regression for granitoid samples RRD10 05 (A) and RRD10 09 (B). Each ellipse represents a single

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116 Figure 3 12. (A) Concordia diagram plotting composite U Pb data for garnet granulite sample RRD10 error. (B) Expanded view of upper end of concordia showing calculated upper intercept age.

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117 Figure 3 13. Concordia diagram plotting composite U Pb discordia r egression for orthogneiss sample LD10 07. Each ellipse represents a single spot analysis

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118 Figure 3 14. (A) Concordia diagram plotting composite U Pb data for orthogneiss sample LD10 08. Each ellipse represents a single spot analysis error. (B) Expanded view of lower end of concordia showing spread of samples and weighted mean age.

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119 Figure 3 15. (A) Concordia diagram plotting composite U Pb data for orthogneiss sample LD10 01. Each ellipse represents a single spot analy error. (B) Expanded view of lower end of concordia showing spread of samples and weighted mean age.

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120 Figure 3 16. Concordia diagram showing U Pb upper intercept regression for quartz pegmatite sample LD10 11. Each ellipse represents a single spot analysis

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121 Figure 3 17. (A) Concordia diagrams plotting composite U Pb data for orthogneiss sample HX (B) Expanded view of upper end of concordia showing samples selected for regression. (C) Expanded view of lower end of concordia showing samples selected for regression.

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122 Figure 3 18. (A) Concordia diagram plotting composite U Pb data for dioritic granulite sample RRD10 20. (B) Concordia diagram plotti ng composite U Pb data for meta granitoid sample LSC10 12. Each ellipse represents a single spot

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123 Figure 3 19. Hf evolution diagram showing the granitoid and meta igneous samples (blue diamonds) against the granitoid samples from the Grassrange (Chapter 2) (orange squares) and data from the Little Belt Mountains (Weiss et al., 2009) (black bars). Lower Crustal lower continental crust calculated using Rudnick and Gao (2003) (dotted line), CHUR chondritic uniform reservoir BSE bulk silicate earth (Faure and Mensing, 2005) (solid line), DM depleted mantle (Chauvel and Blichert Toft, 2001) (dashed line).

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124 Figure 3 20. (A) Probability densi 07. (B) Probability density plot showing Hf T (DM) values for sample LD10 07. Met. metamorphic zircons; Mag. magmatic zircons (see text for more details).

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125 Figure 3 21. Sm Nd evolution diagram sho wing the meta igneous samples (blue diamonds), Little Belt Mountains (green bar), and four different northern Wyoming province crustal outcrops (all citations are in the text). The xenoliths from the Grassrange are shown for reference (orange squares). L ower Crustal zone of lower continental crust calculated using Rudnick and Gao (2003) (dotted line), CHUR chondritic uniform reservoir, BSE bulk silicate earth (Faure and Mensing, 2005) (solid line), DM depleted mantle (Chauvel and Blichert Toft, 20 01) (dashed line), LBM Little Belt Mountains, TR Teton Range, BTM Beartooth Mountains, MR Madison Range, TRM Tobacco Root Mountains.

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126 Figure 3 22. Histogram showing Sm Nd depleted mantle model ages (Ga) of meta igneous whole rocks. T DM calcu lated using the model of DePaolo, 1981.

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127 Figure 3 23. Trace element discrimination diagrams (after Pearce et al., 1984) for meta igneous samples (blue diamonds): (A) Heavy rare earth element Y ( ppm ) vs. high field strength element Nb ( ppm ) (B) H eavy r are earth element Yb ( ppm ) vs. high field strength element Ta ( ppm )

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128 Figure 3 24. Published Archean to earliest Paleoproterozoic U Pb ages of zircon from the northern Wyoming craton (red bars) compared to published ages for the MHB (blue bars) (all ci tations are in the text). OC Owl Creek Mountains, BM Bighorn Mountains, BT Beartooth Mountains, TR Tobacco Root Mountains, SG Sweetgrass Hills, BH borehole. Wyoming light gray, Medicine Hat dark gray. Crystallization ages of meta igneous xenoliths from this study green stars.

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129 Figure 3 25. Comparison of published U Pb ages from the southern Trans Hudson Orogen (all citations are in the text) against the northern Trans Hudson Orogen, Great Falls tectonic zone (Chapter 2 and this study ), and Yavapai Central Plains

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130 CHAPTER 4 LITTLE ROCKY MOUNTAINS Introduction The Archean Medicine Hat block (MHB) lies within a 700 km wide zone of Archean and Paleoproterozoic tectonic elements that separate the Wyoming Craton from the Hearne Province. It is bracketed by the Paleoproterozoic suture Great Falls Mueller et al., 2005; Foster et al., 2006), and the aeromagnetically defined Vulcan structure to the north (Thomas et al., 1987; Hoffman, 1988; Hoffman, 1990; Ross et al., 1991; Eaton et al., 1999) ( Figure 4 1). This zone includes features identified from geophysical surveys (Lemieux et al., 2000), borehole data (Villeneuve et al., 1993), and limited xenoliths (Davis et al., 1995; Gorman et al., 2002). The M edicine H at B lock lies beneath the Phanerozoic sediments of the Western Canada Sedimentary Basin with no reported surface exposures. Because of this, the origin of the MHB and its tectonic relationships to the He arne and Wyoming provinces are largely unknown. The Little Rocky Mountains, Montana, contain excellent sections of Paleozoic and Mesozoic sedimentary rocks (Knetchel 1959) exposed in a domal uplift formed when Tertiary syenite porphyries intruded the sec tion (Weed and Pirsson, 1898). Precambrian basement rocks are also exposed in the Little Rocky Mountains (LRM), but as blocks within and along the margins of the intrusions. These blocks provide one of only two outcrop accesses to otherwise buried baseme nt in the central GFTZ. Southwest of the LRM, the Little Belt Mountains (LBM) provide the only other outcrop access to Precambrian crust within the central GFTZ. The L ittle B elt M ountains were determined by Mueller et al. (2002) via geochronologic and ge ochemical data from

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131 gneisses to be a subduction generated igneous arc formed during the time interval of 1.9 Ga to 1.8 Ga. Based on aeromagnetic data (Sims et al., 2004) and their geographic location within the GFTZ ( Figure 4 2), the Little Rocky Mountain s may comprise a segment of Medicine Hat Block crust, or may represent another opportunity, like the LBM, to gain greater insight into the formation of the GFTZ. This paper presents the results of a geochemical and geochronologic study of the crystalline basement of the Little Rocky Mountains and compares it to the other basement exposure within the central GFTZ, the Little Belt Mountains. We then assess the relationship of the LRM to the Archean Medicine Hat block and Wyoming Province. Whole rock major and trace element geochemistry, whole rock Sm Nd and Pb isotopes, and zircon U Pb geochronology and Hf isotope studies were conducted on 12 orthogneiss, 9 amphibolite, 10 schist, and 3 paragneiss samples. Neoarchean ages from the LRM (this study) and the Priest River complex (Doughty et al., 1998) ( Figure 4 1) are then combined with previously published ages from the MHB (Villeneuve et al., 1993; Davis et al., 1995; Gorman et al., 2002) to compare the MHB with both the Wyoming and Hearne Cratons. Geologic Background The Little Rocky Mountains, located in Philips and Blaine counties of north central Montana, are a series of peaks and buttes exposing Precambrian through Cenozoic rocks ( Figure 4 3). The core of the mountain range consists of Paleogene syenit e intrusions surrounded by a structural dome of basement and cover rocks. Weed and Pirsson (1898) published the first detailed evaluation of the regional geology, observing how the uplift of the Little Rocky Mountains deformed the otherwise flat laying

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132 se dimentary strata of the region, far from the main front of the Northern Rocky Mountains. Stratigraphy The central domain of the Little Rocky Mountains uplift is a series of porphyritic quartz syenite, hornblende syenite, and monzonite intrusions (Bailey, 1974; Russell, 1984). The syenite porphyry includes 2 20 mm diameter phenocrysts of orthoclase, often with albitic plagioclase inclusions, in a fine groundmass of feldspars and quartz (Wilson and Kyser, 1988). Biotite and amphibole occur as rare and gen erally altered mafic phases compr silicified, and host gold and silver mineralization in the region (Peterman, 1980). K Ar ages determined by Hearn et al. (1977) range between 58 and 66 Ma, interpreted as the age of the syenitic intrusion. Adjacent country rock lacks evidence of contact metamorphism, with contacts between the porphyry and country rock interpreted as faults (Wilson and Kyser, 1988). Phanerozoic sedimentary strata are exposed in much of the uplift, includi ng over 1000 m of Paleozoic and 1200 m of Mesozoic strata. The oldest unit in the LRM is the Cambrian Flathead Sandstone, deposited unconformably above pre Belt Supergroup metamorphic units (Knetchel, 1959). The Flathead in turn is depositionally overlai n by the mixed fine grained clastics and carbonates of the Late Cambrian Emerson Formation (Knetchel, 1959). The late Ordovician Bighorn dolomite occurs disconformably above the Emerson Formation, and is a resistant, ridge forming unit locally rich in mar ine invertebrate fossils (Knetchel, 1959). Further up section, above another disconformity, lie Devonian sedimentary units, which are dominantly interbedded carbonate and shale (Knetchel 1959). Carbonates of the Madison group

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133 dominate the Carboniferous s ection, and often form prominent ridges and buttes in the Little Rocky Mountains (Knetchel, 1959). The Mesozoic section includes both marine and non marine sediments, and is dominated by fine grained clastics with subordinate limestone. These include car bonates of the Rierdron Formation, and marine to marine terrestrial shale of the Swift and Morrison formations (Knetchel, 1959). Lower Cretaceous deposition is represented by the sandstone and shale of the Kootenai formation. This is overlain by the shal e rich Thermopolis, Mowry, and Warm Creek formations (Knetchel, 1959). The Montana group concludes the Cretaceous sequence, with both sandstone and shale formations. The Bearpaw Formation of the Montana Group is the youngest Mesozoic unit in the Little R ocky Mountains, consisting of marine shale with minor bentonite and cherty horizons (Knetchel, 1959). Precambrian gneiss, schist, and amphibolite crop out in the Little Rocky Mountains in stream gorges and saddles between topographic highs (Weed and Pirs son 1896). Knetchel (1959) observed that many of the Archean metamorphic rocks are foliated meta sedimentary rocks and meta volcanic rocks, based on mineralogy and stratigraphic relationships between biotite rich gneiss and schist units, quartzites, and a mphibolites. The metasedimentary rocks are dominantly quartzofeldspathic, with variable contents of garnet, chlorite, biotite and white mica, and amphibole (Knetchel, 1959). In the Ruby Gulch area, lithologies include fine grained amphibolite, finely ban ded white gneiss, sheared gneiss with elongate feldspars, a variety of micaceous, garnet bearing, and amphibole schists, and thin beds of quartzite (Dyson, 1939). Peterman (1980) collected seven samples from road cuts near the Ruby Gulch mine site. These included several samples of fine grained biotite granitic gneiss, two samples

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134 of biotite hornblende tonalitic granodioritic gneiss, and one fine grained quartz amphibolite. Biotite in these samples is relatively fresh, and plagioclase grains display moder ate sericitization (Peterman, 1980). The samples displayed strong fabrics, dominated by biotite foliation and (in some samples) quartz feldspar augen (Peterman, 1980). Previous Geochronology A nd Geochemistry Limited age control exists for the Precambrian units of the Little Rocky Mountains. K Ar dates published by Burwash et al. (1962) yielded ages of 1.71 Ga (hornblende) and 1.75 Ga (biotite), consistent with other cooling ages for the region (Giletti, 1966; Roberts et al., 2002). Rb Sr isotopic analyse s by Peterman (1980) yield a seven point isochron corresponding to an age of ~2.55 Ga, and an initial 87 Sr/ 86 Sr value of ~0.706. Sample Descriptions The samples collected for this study fall into several broad categories: Amphibolite, meta supracrustal sc hist and gneiss, and orthogneiss. These are consistent with previous observations by Weed and Pirsson (1896) and Knetchal (1959). Sample mineralogy and characteristics are discussed below. Amphibolite Samples MLR 03 ( Figure 4 4A, B), MLR 07 and MLR 14 ar e garnet amphibolites. They are dominated by fresh, anhedral to subhedral green amphibole Garnet occurs as irregular fragmented masses with quartz, feldspar, and opaque inclusions. Trace volumes of opaque phases are present. MLR 10, MLR 11, and

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135 LRMGM 2 ( Figure 4 4C, D) are medium to fine grained garnet quartz amphibolites. Amphibole grain or ientation and limited compositional variation between quartz poor and rich layers define a weak fabric in the samples. Amphibole is green to brownish green, with anhedral grain shapes and embayed grain boundaries common. Grain size ranges between 0.1 to 0.5 mm in width. Plagioclase is the second most abundant phase, occurring as anhedral grains with lobate margins. Trace amounts of quartz grains are present and exhibit undulose extinction. Garnet occurs as subhedral grains up to 1 mm in width. The gra ins are fresh and contain abundant plagioclase and uncommon quartz inclusions. LRMGM 1, LRGM 6, and LRMPG 1 are medium to fine grained quartz amphibolites. Green to brown green amphibole grains, generally between 0.1 0.5 mm wide and up to 1 mm long, domi nate the samples. Amphibole grain orientation defines a lineation in the samples, and the grains are generally subhedral, with some embayed grain margins. Smaller anhedral plagioclase grains with generally lobate grain boundaries make up ~30 40% of each sample. The remainder is uncommon subhedral to anhedral quartz. Quartzofeldspathic Schists Sample LRMGM 3 and MLR 08 are medium grained, amphibole biotite bearing quartzofeldspathic schists. A foliation is defined by biotite, with compositional variatio n and amphibole orientation contributing to the overall fabric. Rounded plagioclase porphyroblasts up to 1 mm in diameter deflect the biotite foliation. Matrix plagioclase grains are smaller (only up to 0.5 mm in diameter), subhedral, and show some seric itic alteration. Uncommon microcline grains are similarly sized and typically less altered than plagioclase. Quartz is elongate, with grains up to 1 mm in length. Dynamic recrystallization of quartz is indicated by subgrain formation and undulose extinc tion.

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136 Biotite grains are up to 0.5 mm in length, and are generally unaltered. Green to green brown amphibole is subhedral, and shows evidence of alteration along cleavage planes. Samples LRMGM 4, LRMGM 5, and LRMPG 3 are similar to LRMGM 3, but include rare, euhedral titanite grains. MLR 16, MLR 18, MLR 18, LRMGM 7 ( Figure 4 5E) and LRMPG 2 ( Figure 4 5F) are fine grained biotite quartzofeldspathic schists. Biotite and a weak compositional variation define the layering within the samples. Biotite grain s are euhedral to subhedral, and up to 0.2 mm in length. Quartz is the dominant mineral, with subhedral grains <0.2 mm in width making up over 50% of the samples. Plagioclase and K feldspar are common, with unaltered subhedral grains similar in size to t he quartz grains. Gneisses Sample LRMGM 8 ( Figure 4 4E, F) is a biotite granitic gneiss with abundant rounded microcline and perthite grains up to 3 mm in diameter. Biotite defines a weak foliation in the sample. Quartz is abundant and shows extensive subgrain development and undulose extinction consistent with dynamic recrystallization. Samples MLR 04 ( Figure 4 5C, D) and MLR 05 are garnet biotite quartzofeldspathic augen gneisses. Plagioclase grains core large porphyroclasts, often >2 mm in diameter A biotite foliation and quartz ribbon lineation define a strong fabric in the sample. The quartz ribbons show extensive polygonal subgrain rotation recrystallization diagnostic of mid temperature (~400 500 C) metamorphic dynamic recrystallization. Relatively unaltered, euhedral garnet porphyroblasts are present in the sample. Sample MLR 13 ( Figure 4 5A, B) is a staurolite chloritoid quartzofeldspathic gneiss, with cm scale compositional variations defined by chloritoid rich and poor

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137 domains. Chloritoid defines the foliation in the sample, and is scattered throughout the rock. Staurolites up to 2 mm in length are common. The grains commonly contain anhedral to subhedral quartz inclusions. Dynamically rec rystallized, elongate quartz grains dominate the rock matrix, in association with uncommon k feldspar. Uncommon euhedral titanite grains occur throughout the chloritoid and staurolite domains of the sample. Samples MLR 06, MLR 09, and MLR 19 are biotite q uartzofeldspathic gneisses. The samples contain biotite rich and k feldspar rich bands, which are several mm thick. The sample fabric is defined by biotite foliation; recrystallized, elongate quartz grains and ribbons; and augen cored by feldspar. Gray gneiss samples MLR 15 and MLR 12 differ somewhat from MLR 06, MLR 09 and MLR 19, with a dominant granoblastic fabric and the presence of amphibole. Other Samples Sample MLR 01 is a garnet, two mica, quartzofeldspathic mylonitic gneiss. Quartz dominates th e sample, and occurs as dynamically recrystallized, elongate grains to ribbons. Microcline grains are up to 0.5 mm in length, blocky, with irregular grain boundaries. Large microcline porphyroclasts are common as well, and deflect quartz ribbons. Biotit e and white mica grains define a weak foliation, and are scattered through the sample. Garnet occurs as euhedral, relatively fresh porphyroblasts with uncommon inclusions of quartz and fractures. Samples LRM 1, LRM 3, LRM 5, LRM 6, and LRM 7 were received from Dr. Zell Peterman as vials of zircon separates having been previously prepared for Rb Sr analyses (Peterman, 1981). Samples LRM 1, LRM 5, LRM 6 and LRM 7 were 3 was

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138 descri (Peterman, 1981). Results U Pb and Lu Hf data from meta plutonic and detrital zircons in addition to whole rock geochemical and isotopic data from samples collected from the Littl e Rocky Mountains, Montana, are presented below. Sample locations are shown in Figure 4 3. Methods are described in Appendix A. Geochemical and isotopic data and the latitude and longitude of the sample locations are summarized in Tables B 7 through B 9 (Appendix B). Whole Rock Geochemistry All of the samples from this study are metamorphic, and when examining the samples in thin section, protolith determination is sometimes uncertain. Representative photomicrographs are shown in Figure 4 4 and 4 5. Based primarily on the diversity of ages of the zircon populations, 11 samples were determined to be meta igneous (see below). The orthogneissic samples (Table 4 1) have silica contents ranging from 55.3 to 73.8 wt. %. Collectively, the orthogneiss sampl es cluster along the metaluminous peraluminous boundary ( Figure 4 6) when plotted according to the alumina saturation index of Shand (1943). Meta igneous amphibolites (Table 4 2) yielded silica contents ranging from 46.5 to 53.1 wt. %. Schist samples ML R 08, MLR 17, and MLR 18 yielded a diverse range of zircon ages that are compatible with a sedimentary protolith. Gneissic samples MLR 04, MLR 05, and MLR 13 also yielded a diverse range of zircon ages, compatible with a sedimentary protolith. The schist s (Table 4 3) primarily fall within the metaluminous field, but with some samples extending into the peraluminous field of the alumina

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139 saturation index of Shand (1943) ( Figure 4 6). The paragneisses (Table 4 4), mostly fall within the peraluminous field o f the alumina saturation index of Shand (1943) ( Figure 4 6). The protolith(s) of these schists and paragneisses may be sedimentary, volcaniclastic, and/or volcanic based on major element compositions. Discrimination is determined by comparing wt. % value s of MgO and CaO versus Al 2 O 3 Weathering processes leave behind residual Al 2 O 3 generally in clays, relative to MgO and CaO. MgO and CaO are carried away as a dissolved load when they are liberated by the breakdown of less stable igneous phases at surfa ce conditions. Leyreloup et al. (1977) developed a ternary diagram, based on CaO MgO Al 2 O 3 for discrimination of igneous and sedimentary protoliths. Meta sedimentary units are recognized by their abundance of Al, which is concentrated preferentially in sediments as plagioclase and mafic mineral weather. Abundances of Al 2 O 3 are relatively low (12.8 to 18.9 wt. %), but alkali contents in some are sufficiently high that their normative mineralogy contains 0 7% corundum. Petrographic discrimination betwe en sedimentary, volcaniclastic, and/or volcanic origins for the schists is difficult; however, variations of Ca, Mg, and Al concentrations provide some insight into their origins. Among the schistose samples, all plot in the meta sedimentary field, with o ne lying on the border of meta igneous, and all three of the paragneiss samples plot within the meta sedimentary field. Although most of the schistose samples have compositional characteristics that indicate a sedimentary origin, these discriminants are n ot sufficiently accurate to exclude a volcaniclastic component. Trace element compositions (Table 4 1 and 4 2) of the orthogneiss and amphibolite samples are summarized in Figure 4 8A, and plotted normalized to primitive mantle values of McDonough and Sun (1995). The plots show enrichments in fluid

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140 mobile incompatible element concentrations, such as Rb, Ba, and Pb, up to 400 times the primitive mantle concentrations. This relative enrichment is paired with minimal enrichment in fluid immobile trace elemen ts, including heavy REE ( Figure 4 8A). The majority of the orthogneiss samples yielded low values of Eu and Sr relative to elements of similar compatibility, while the amphibolite samples yielded higher values of Eu and Sr relative to elements of similar compatibility ( Figure 4 9). Nb and Ta are depleted relative to the observed values for neighboring elements in all of the orthogneiss and amphibolite samples when normalized to primitive mantle. U Pb Geochronology O f Zircon Rock samples larger than ~20 cm in diameter were processed for zircon separates in addition to whole rock geochemistry. The zircon concentrates were then analyzed for U Pb ages and Hf isotopic composition to aid in characterizing the events recorded in the LRM. Zircon U/Pb age populat ions from meta supracrustal samples were filtere geochronology community, and served to filter out grossly disturbed grains from the sample set. Zircon failing this test typically represents disturbed grains, where Pb loss occ urred during some past metamorphic or fluid flow event. Some analyses are mean and standard deviation of the main age population. Grains passing through the filters typ ically define a discordia line, and are used to generate upper intercept ages. Crystallization ages for the meta plutonic samples (orthogneisses and amphibolites) discussions b generate age distributions for each meta supracrustal sample.

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141 U Pb zircon ages from orthogneiss samples LRM 3, LRM 5, LRM 6, LRM 7, MLR 01, MLR 06, MLR 09, MLR 15, and MLR 19 yielded a wide spread of ages, from 1.87 Ga to 3.29 Ga, despite the relatively uniform bulk chemistries (Table 4 5). Zircons from 8 of the 9 samples reveal 2 distinct populations and a single sample of 207 Pb/ 206 Pb ages, and a third age is represented by a single sample. The first age population includes samples MLR 01 and MLR 06. MLR with a scatter of ages from ~2.39 to 2.44 Ga ( Figure 4 10A). The scatter within the sample indicates that the isotopic system has likely been disturbed. The weighted mean 207 Pb/ 206 Pb age for the entire data set is 2.42 0.01 Ga, which is interpreted as a minimum crystallization age. MLR scatter of ages from ~2.38 to 2.45 Ga ( Figure 4 10B). Similar to MLR 01, the scatter indicates that the isotopic system was l ikely disturbed. The weighted mean 207 Pb/ 206 Pb age for the entire data set is 2.42 0.01 Ga, which is interpreted as a minimum estimate of the crystallization age. The second orthogneiss age population includes samples MLR 15, MLR 19, LRM 3, LRM 5, LRM 6, and LRM 7, which yielded crystallization ages of ~2.8 Ga. MLR 15 and MLR 19 yielded ages of 2.78 0.01 Ga, and 2.79 0.01 Ga respectively, with no inheritance of older grains ( Figure 4 11). MLR discordant), 16 of which were excluded from the crystallization calculation due to large measurement error (>3% Ma error), or due to anomalously young ages. These young ages reflect Pb loss, causing grains to trend along a d iscordia line away from the main 15 yielded an age of 1.87 0.07 Ga, which is interpreted as a metamorphic age. MLR 19 yielded 56 grains

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142 are attributed to Pb loss causing grains to trend along a discordia line away from the main population. With these 6 grains excluded, 50 grains went into the crystallization calculation. LRM 3 yielded 23 zircons, 13 of which appear to fall along a disco rdia line ( Figure 4 12A). The oldest 6 grains yielded an upper intercept age of 2.80 0.09 Ga ( Figure 4 12B). LRM 5 yielded 55 zircons, only 8 of which had younger ages and higher discordance between 7 10% and 4 had older ages and Pb loss, and these grains were excluded from the c ry stallization calculation. The four Figure 4 13). LRM 6 yielded 6 zircons total, with only 2 meeting the concordance criteria 6 did not yield a crysta llization age as any calculation with so few analyses is of dubious reliability. LRM discordant. Six grains went into the upper intercept age calculation of 2.77 0.04 Ga ( Figure 4 14). Two grains were excluded becaus e of high measurement errors, and 4 were excluded for being outliers. Orthogneiss sample MLR 09 is the only orthogneiss sample that yielded a were excluded for being outlie rs. When all of the grains are plotted on concordia ( Figure 4 15A) analyses scatter between ~3.2 and ~2.8 Ga indicating that metamorphism at c. 2.8 Ga likely disturbed the isotopic system. Ten analyses were rejected due to possible Pb loss shifting their ages younger. The remaining 10 grains

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143 yielded a weighted average 207 Pb/ 206 Pb age of 3.20 0.01 Ga ( Figure 4 15B) which is interpreted as a minimum crystallization age. Amphibolites MLR 03 and MLR 10 were large enough for zircon separation. MLR 03 yield ed 58 grains zircons are plotted on concordia ( Figure 4 16A) analyses scatter between ~3.0 and ~2.8 Ga indicating that metamorphism likely disturbed the isotopic system similar to MLR 09. 20 analyses were reje cted due to significant Pb loss. The remaining 12 grains yielded a weighted average 207 Pb/ 206 Pb age of 3.01 0.01 Ga ( Figure 4 16) which is interpreted as a minimum crystallization age. Sample MLR 10 yielded 60 zircons total, with only 5 low discordance 207 Pb/ 206 Pb ages between 3.10 Ga and 3.17 Ga This hindered the calculation of a concordia intercept age, and indicates that the sample likely records multiple Pb loss events. The schist and paragneiss samples yielded 1 3 2 detrital zircon grains. The three schist samples (MLR 08, MLR 17, and MLR 18) yielded zircons ranging from sub rounded to well rounded. In aggregate, 207 Pb/ 206 discordant range from ~2.57 Ga to ~3.32 Ga (Table 4 6). On a p robability density plot there are prominent age peaks at ~2.79 Ga (65% of the grains, from 2.74 2.81 Ga) and ~3.19 Ga (5% of the grains, from 3.13 3.20 Ga) ( Figure 4 17). The three paragneiss samples (MLR 04, MLR 05, and MLR 13) yielded zircons rangin g from subhedral to well rounded. 207 Pb/ 206 from ~1.78 Ga to ~3.20 Ga (Table 4 6). The data have a prominent peak at ~2.79 Ga (41% of the grains, from 2.76 2.80 Ga) ( Figure 4 17). Figure 4 17 shows a proba bility

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144 distribution function comparing detrital zircon age spectra from the schists and paragneisses; data from the orthogneiss samples are provided for comparison. Hf I sotopes I n Zircon Zircons were chosen for Lu Hf analysis based on a number of differ ent factors. Samples were arranged by age to better select a representative population of ages for Hf analysis. Zircons were preferentially chosen for low discordance (as discussed above) to have a more accurate U Pb age with which to reduce the Hf isotop e data. Zircons were further chosen for zonation patterns and grain size. If a zircon showed zonation and a U Pb analysis was undertaken in a particular zone, ideally, the Hf spot would also be analyzed within the same zonation. The Hf spot size is 40m and the (initial) calculations used rock crystallization ages for meta plutonic samples, and 207 Pb/ 206 Pb ages of individual detrital grains. For mantle separation ages, Hf T DM represents minima because Lu/Hf in zircon is invariably lower than in whole rocks (e.g., Griffin et al., 2002). Calculation of 2 stage or crustal residence ages, however, requires knowledge of the Lu/Hf of source(s) and were not calculated because the isotopic compositions are likely to result from a mixing of crustal and mantle sources. For those meta plutonic samples (orthogneisses and amphibolite) that had an estimated time of pro calculated in two different ways. The first was based on the individual age (IA) of each analyzed grain within each sample regardless of discordance or Pb loss. This range of 5 labeled as (IA) (IA) ~3.0, or ~3.2 Ga) for each sample. This range of values was also recorded in Table 4 5 (T) (T)

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145 standard deviation used to characterize the spread of values (Table 4 are plotted on Figure 4 18A vs. 207 Pb/ 206 Pb age (Ma). The youngest group of orthogneiss samples (MLR 01 and MLR 06) yiel ded U Pb (2.4 Ga) for these samples ranges from 2.2 to 11.6 (Table 4 (IA) values for the same samples ranged from 2.7 to 16.1 (Table 4 5). These differences likely reflect variable input of juvenile mantle derived versus evolved crustal materials in the genesis of the parental magma over the course of zircon crystallization. The more coherent Hf isotopic compositions calculated at ~2.4 Ga suggest that the younger ages reflect Pb loss and that the Lu Hf system remai ned largely closed during a younger metamorphic event. These samples give a limited range of average depleted mantle model ages (T DM ) from 2.97 Ga to 3.00 Ga with standard deviations of 0.04 and 0.10 (Table 4 5). (2.8) values for orthogneiss samples MLR 15, MLR 19, LRM 3, LRM 5, and LRM 7 were also calculated (Table 4 (2.8 Ga) zircons from the samples range from 6.6 to 14.5, averages of 9.4 to 12.9, and standard deviations of 0.8 to 1.9 (Table 4 5). Because LRM 6 did not yield a (T) values were not calculated. Average Hf T (DM) for the ~2.8 Ga orthogneiss samples range from 3.31 Ga to 3.53 Ga with standard deviations from 0.03 to 0.07 (Table 4 5). The oldest orthogneiss sample (MLR 09) yielded a U Pb (3.2 Ga) 09 range from 3.8 to 2.2 with an average of 0.1 and a standard deviation o f 1.7. As with the younger samples, calculating the initial Hf (T) ) yields more coherent estimates of initial Hf compositions than using the 207 Pb/ 206 (IA) ), i.e., the

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146 discordant age s represent Pb loss rather than extraneous grains. Open system behavior, however, cannot be excluded. The amphibolite samples (MLR 03 and MLR 10) yielded some zircons large enough for Hf isotopes to be analyzed. In sample MLR 03, the measured Hf isotopic compositions for all zircons were recalculated to the U Pb crystallization age of 3.01 Ga. (3.0 Ga) 6.1 with an average of s calculated for the individual age (IA) of each analyzed grain (~2.7 to ~3.0 Ga) within MLR 03, the ranged from 3.6 to 9.8 (Table 4 5). The average Hf T (DM) for sample MLR 03 was calculated to be 3.28 Ga with a standard deviation of 0 .09. MLR 10 did within MLR ranged from 2.4 to 4.4 (Table 4 5). The average Hf T (DM) for sample MLR 10 was calculated to be 3.40 Ga with a standard deviation of 0.08. All three schist samples yielded zircons large enough for paired U Pb and Lu Hf isotopic analysis. Initial Hf isotopic compositions wer e calculated using the 207 Pb/ 206 Pb age of each individual zircon. MLR Ga to 15.7 at 2.48 Ga, and Hf T (DM) ranging from 3.27 Ga to 3.44 Ga (Table 4 6). MLR 6.9 at 2. 82 Ga to 18.5 at 2.57 Ga, and Hf T (DM) ranging from 3.32 Ga to 3.69 Ga (Table 4 6). MLR 18 8.5 at 2.80 Ga to 16.0 at 2.78 Ga, and Hf T (DM) ranging from 3.44 Ga to 3.63 Ga (Table 4 6)

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147 All three paragneiss samples yielded zircons large enough for U Pb and Hf isotopic analysis. Initial Hf isotopic compositions were calculated using the 207 Pb/ 206 Pb age of each individual zircon. MLR Ga to 16.0 at 1.78 Ga, and Hf T (DM) model ages ranging from 2.57 Ga to 3.01 Ga (Table 4 6). MLR from 1.1 at 3.20 Ga to 7.1 at 2.37 Ga, and Hf T (DM) model ages ranging from 2.75 Ga to 3 .44 Ga (Table 4 6). MLR that range from 7.7 at 2.78 Ga to 13.3 at 2.64 Ga, and Hf T (DM) model ages ranging from 3.32 Ga to 3.53 Ga (Table 4 6). Sm Nd Whole Rock Isotopes As a group, the orthogn eiss samples (including those without U Pb ages) show a range in whole (0) from 20.2 to 50.2. Initial ratios were calculated for samples MLR 01 and MLR 06 using the best estimates of the individual crystallization ages shown in Figure 4 10 ( 207 Pb / 206 (2.4) of 5.7 and 7.3 (Table 4 7, Figure 4 18B). Values for samples MLR 15, MLR 19, LRM 3, LRM 5, and LRM 7 were calculated using the 207 Pb/ 206 Pb ages shown in Figure 4 11 to Figure 4 14 (i.e., ~2.8 Ga), which yie (2.8) of 6.5 to 10.5 (Table 4 7, Figure 4 18B). 09 was calculated using the crystallization age of 3.19 Ga ( Figure 4 (3.2 Ga) of 0.9 (Table 4 7, Figure 3 18B). Sm Nd depleted mantle model ages were calculated for each sample using the model of DePaolo (1981). Orthogneiss samples MLR 01 and MLR 06 yielded Sm Nd T (DM) calculated to be 3.55 Ga and 3.15 Ga respectively ( Figure 4 10). Samples MLR 15, MLR 19, LRM 3, LRM 5, and LRM 7 yield ed a range of Sm Nd T (DM) model ages from 3.28 to 3.80 Ga (Table 4 7). Sample MLR 09 yielded a Sm Nd T (DM) of 3.42 Ga (Table

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148 4 7). The amphibolite samples (including those without U Pb ages) show a range in (0) values from 3.8 to 21.0. Initial ratios were calculated for sample MLR 03 using the best estimate of crystallization age ( 207 Pb/ 206 Pb age 3.01 Ga) shown in Figure 4 16. (3.0) of 2.0 (Table 4 7, Figure 4 18B). The schist samples show a (0) values from 30.1 to 50.7, and Sm Nd T (DM) ranging from 2.99 Ga to 3.72 Ga (Table 4 7). Discussion Geochemical I nsight I nto S ample O rigins Geochemical evidence suggests that the orthogneiss samples either formed in a subduction zone environment, or represent remobilization of rocks that have an arc signature. Most of the meta plutonic are metaluminous, and one is peraluminous ( Figure 4 6), which is not particularly diagnostic, but is consistent with observed granito id suites from continental arc and continental collision settings (e.g., Mainar and Piccoli, 1989; Rogers and Hawkesworth, 1989; Chappell and White, 2001; Villaseca et al., 2012). Trace element patterns observed in the orthogneiss samples are also consist ent with patterns in rocks from modern continental arcs ( Figure 4 8A; e.g., Pearce, 1983; Thompson et al., 1984). Similarities include negative Eu and Sr anomalies indicative of fractional crystallization or residual plagioclase in the source, negative Nb and Ta anomalies, and elevated concentrations of some fluid mobile elements such as Ba and Pb. Chondrite normalized REE plots reveal elevated concentrations of light versus heavy rare earth elements, with non enrichment to depletion in heavy REE suggesti ve of residual garnet and/or amphibole pyroxene in the source region (La/Yb values between 150 and 1.5; Figure 4 9).

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149 The meta supracrustal samples are difficult to characterize in terms of protolith. Trace element patterns (Fig 4 8B) broadly mirror those of the meta plutonic sample suite (Fig 4 8A), and again are consistent with a convergent margin provenance (Pearce, 1984; Rogers and Hawkesworth, 1989; Taylor and McLennan, 1995; Chappell and White, 2001). When the samples are plotted on a ternary diagram of the CaO MgO Al 2 O 3 system ( Figure 4 7) (Leyreloup et al., 1977), most of the schist samples lie Little Rocky Mountains as Exposed Medicine Hat Block Crust Correlation of igneous ages Due to the lack of exposure of Medicine Hat Block crust, it remains poorly characterized. Data presented in this study link the Little Rocky Mountains to the Medicine Hat Block, and provide a valuable opportunity to characterize materials otherwise known mostly from geophysical evidence. Xenoliths from the Sweetgrass Hills yield zircon U Pb ages of 2.60 to 2.84 Ga Davis et al., 1995; Gorman et al., 2002). Similarly, Villeneuve et al. (1993) analyzed basement samples from core penetrating the MHB crust, revealing U Pb zircon crystallization ages of 2.62 to 2.72 Ga, and one crystallization age of 3.28 Ga. Orthogneiss samples from this study reveal three distinct crystallization age ranges in the LRM: 2.42 Ga for two of the nine orthogneiss es, 2.77 Ga to 2.81 Ga for five of the nine samples, and finally one sample at 3.19 Ga ( Figure 4 19). The ninth orthogneiss sample did not yield a crystallization age. These ages overlap with the documented MHB ages noted above, suggesting a geochronolog ic link between the blocks. The abundance of Neoarchean crystallization ages in the LRM suggests that they share a closer affinity with the MHB than the Wyoming craton, which preserves an extensive Mesoarchean igneous history. Mueller et al. (1988; 1993; 2010)

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150 report extensive c. 2.8 Ga magmatism in the Beartooth and Madison ranges. In contrast to the current MHB data, however, the Wyoming Craton also contains abundant Paleoarchean ages ( Figure 4 19). Mogk et al. (199 0 ) and Mueller et al. (1993) report abundant TTG gneiss units in the Madison Range with c. 3.1 to 3.3 Ga crystallization ages. Data from the Tobacco Root Mountains reveals this older age component as well, Krogh et al. (2011) dating tonalitic gneiss and amphibolite samples at c.3.33 Ga and c. 3.34 Ga respectively. While age matching is by no means conclusive, the lack of igneous ages >3.2 Ga in the Little Rocky Mountains ( Figure 4 19) suggests a closer tie to the MHB than the Wyoming Craton. Whole rock Sm Nd and zircon Hf isotopic data ca n also be used in evaluating the crustal affinities of the LRM, relative to the Wyoming Craton in light of the absence of MHB isotopic data. Three of the orthogneiss samples (c. 2.4 Ga and c. 3.2 Ga) and the amphibolite sample (c. 3.0 Ga) do not have a di rect age correlation with the northern Wyoming craton ( Figure 4 18A). The c. 2.8 Ga orthogneiss samples are of a similar age to the Beartooth Mountains. The Beartooth Mountain Nd (2.8) values range from 0.3 to 6.0, while the LRM orthogneiss samples have much lower Nd (2.8) values, between 6.5 to 10.5. This indicates that formation of the LRM ~2.8 Ga orthogneiss involved a larger percentage of evolved (crustal) material than rocks of the same age from the Beartooth Mountains in the Wyoming craton. Insi ghts into meta sediment provenance Detrital zircons extracted from presumed meta supracrustal units of the LRM reveal a dominance of Neoarchean ages, in line with published Medicine Hat Block magmatic events ( Figure 4 17). Zircon 207 Pb/ 206 Pb ages from the schist samples (disco

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151 ages based on the youngest zircons within each sample are variable, ranging from 2.57 to 2.78 Ga. Two of the 3 samples (MLR 17, MLR 18) show a prominent age peak at ~2.8 Ga, and the final sample (MLR 08) contains a small population of zircons of that age. This age overlaps with the crystallization ages of 6 of the orthogneiss samples from the LRM, as discussed above. The ~2.8 Ga detrital grains from the schist samples ( Figure 4 18A 6.9 to 16.6, indicating derivation from an evolved source region (Table 4 (c. 2.8 Ga) values overlap with initial 6 .6 to 14.5, and suggest that the protolith sediments may be locally derived. However, there is a large amount of ~2.8 Ga in the Beartooth Mountains from the Wyoming craton (0) for the schist samples is relat ively large, from 36.3 to 50.7, and the Nd depleted mantle model ages range from 3.47 Ga to 3.72 Ga (Table 4 7). The model ages for these meta sediments suggest that they received detritus from older crust, possibly reflecting a contribution from the ol der material documented in the MHB. The Sm Nd isotopic data corroborate the zircon Hf data, indicating significant contributions from evolved materials into the protolith sediments. Paragneiss MLR 13 yielded a similar age range to the schist samples dis cussed above, from ~2.64 Ga to ~2.82 Ga with a prominent age peak of ~2.80 Ga ( Figure 4 17). The ~2.8 Ga detrital grains from MLR 13 ( Figure 4 ranging from 7.7 to 12.9 (Table 4 rital from 6.6 to

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152 contributed to the petrogenesis of their source. In contrast to MLR 13, s amples MLR 04 and MLR 05 yielded few low discordance zircons and are dominated by younger Neoarchean ages ( Figure 4 17). MLR but they range from 1.78 Ga (one grain) to the other 3 between ~2.55 Ga and ~2.59 Ga. Given the tectonic and magmatic activity in the GFTZ during the Paleoproterozoic (see Chapters 2 and 3), the single zircon grain from MLR 04, which yielded the 1.78 Ga age, is likely to be metamorphic. MLR 0%, which range from ~2.29 Ga to ~2.61 Ga. A single Paleoarchean grain yields a 207 Pb/ 206 Pb age of 3.20 Ga. While a single detrital zircon age is of limited significance, this age does correlate with data from Villeneuve et al. (1993) documenting c. 3.2 Ga material in the MHB. (0) values for the paragneiss samples are 30.1, 31.2, and 43.6 for MLR 04, MLR 05, and MLR 13, respectively (Table 4 7). The Sm Nd depleted mantle model ages are 2.99 Ga, 3.04 Ga, and 3.69 Ga respectively (Table 4 7). Both the Hf in zircon and Sm Nd in whole rock systematics indicate a mixture of sources in the provenance of these samples. Sm Nd data from MLR 13 support the Hf in zircon data of a greater evolved crustal contribution, whereas the MLR 04 and MLR 05 seem to have a somewhat greater juvenile component. Both age data and Hf isotopic data are consistent with a MHB provenance for the meta supracrustal samples. Meta supracrustal units deposited on the northern margin of the Wyoming Craton (representing the Beartooth, Ruby, and Tobacco R oot mountains) between 2.7 Ga and 3.2 Ga would likely yield age spectra far richer in 3.2 Ga 3.4 Ga zircons (Mueller et al., 1998). Source areas for these ages are particularly

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153 well represented in the Madison and Tobacco Root mountains, where extensive TTG magmatism occurred between 3.1 Ga to 3.5 Ga (Mogk et al. 199 0 ; Mueller et al., 1993; Krogh et al., 2011). Although data from unequivocal MHB is limited, the oldest age occurs at c. 3.27 Ga, and is documented from a single drill core sample (Ross et al ., 1991). Taken in conjunction with the younger igneous age ranges observed from the LRM samples, the detrital zircon data favor a MHB affinity over a Wyoming Craton provenance for the meta supracrustal rocks. Other regional considerations The Little Rock y Mountains and the Little Belt Mountains to the southwest represent the most extensive basement exposures in the GFTZ; however, they differ greatly in character. The Little Belt Mountains (LBM) are composed of predominantly dioritic to granitic gneisses and migmatites (Pirsson, 1900; Weed, 1900; Schafer, 1935; Vogl et al., 2003). Zircons analyzed from three of the dioritic rocks, as well as trace elements and Nd isotopes led Mueller et al. (2002) to the conclusion that the rocks exposed in the LBM develo ped at ~1.86 Ga in response to subduction between the MHB and the Wyoming craton. Similar to the LBM, trace elements from the Little Rocky Mountains also are similar to those of modern convergent margins with classic depletions in high field strength elem ents (HFSE) relative to the large ion lithophile elements (LIL) (e.g., Thompson et al., 1984). However, unlike the Paleoproterozoic Little Belt Mountains, the Little Rocky Mountains appear to be primarily Archean in age, with geochronologic results that a re quite varied. Orthogneiss and amphibolite samples range from ~2.4 Ga up to ~3.2 Ga, and detrital grains from the meta supracrustal schists and paragneisses ranging from ~1.8 Ga up to ~3.3 Ga. The whole rock Nd (Mueller et al., 2002) and zircon Hf (Wei ss et al., 2006) isotopic data from the LBM both

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154 suggest extensive involvement of the depleted mantle in the petrogenesis of the samples (Mueller et al., 2002; Chapter 2 and 3). The ~3.2 Ga orthogneiss and ~3.0 Ga amphibolite from the LRM show similar pat terns, but the ~2.8 Ga orthogneisses from the LRM have Hf and Nd data that lie below CHUR/BSE and within or just above zones with slopes indicative of lower crustal values ( Figure 4 18). The Little Belt Mountains appear to have formed as part of the GFTZ in an ocean subduction convergent margin environment, and the Little Rocky Mountains represent older earliest Paleoproterozoic to Archean continental crust. Doughty et al. (1998) was the first to date the Archean basement of the Priest River complex ( Figur e 4 1), and documented an age of 2.65 Ga from U Pb analyses of zircon. This was suggested to indicate that the Priest River Complex is either a western extension of the Rae Hearne craton, or that it is a unique block separated from North American Archean crust by a domain of Early Proterozoic crust. Similarly, the Clearwater metamorphic core complex, located south of the Priest River complex and west of the MHB, also yields an Archean basement age (~2.67 Ga) from gneiss (Brewer et al., 2008; Jansen et al. 2011). Guevera et al. (2012) and Jansen et al. (2011) also report 1.87 Ga ages from orthogneiss samples within the Clearwater complex. These ages are similar to those found in orthogneiss samples from the LBM along the southern margin of the Archean MH B. Foster et al. (2006) suggested that, based on the Archean ages (Doughty et al., 1998) and inferred by isotopic tracer data and xenocrystic zircons from Eocene plutons (Whitehouse et al., 1992), the Archean basement in the Priest River and Clearwater co mplexes are instead an extension of the Medicine Hat block.

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155 Crystallization ages from the Rae Hearne craton range from ~2.54 to ~2.73 Ga, with major plutonic events occurring between 2.58 and 2.57 Ga ( Ray and Wanless, 1980 ; Bickford et al., 1987 ; Annesley et al., 1992 ; Bickford et al., 200 1 ). The oldest known components of the Peter Lake domain, which represents the southeastern component of the Hearne craton, are tonalitic to granodioritic gneisses >2.58 Ga ( Rayner et al., 2005). Within the age span of ~ 2.54 to ~2.73, Hearne ages correlate well with MHB and LRM ages. However, Hearne is lacking in ages approximating the ~2.8 Ga ages that are seen in the MHB and LRM. Conclusions Igneous crystallization and detrital zircon age correlation are consistent w ith the L ittle R ocky M ountains sharing a strong tie to the M edicine H at B lock rather than with the generally older Wyoming province. As such, the Archean basement of the Little Rocky Mountains likely represents the only surface exposure of Medicine Hat b lock crust. U Pb data for meta plutonic samples taken from the LRM range from 2.42 to 3.19 Ga, documenting numerous crustal building episodes in the LRM. Detrital zircon ages range from 1.78 to 3.32 Ga and are particularly abundant at c. 2.8 Ga, recordin g what is likely the period of most active crustal growth in the LRM. Geochemical data from the Little Rocky Mountains show H FSE depletions and LIL enrichments characteristic of subduction zone magmatism. The limited enrichment of HREE relative to primit ive mantle values reflects the source mineralogy, likely lower crustal materials where garnet is residual after melt extraction. This pattern suggests partial melting deep enough for garnet to remain stable, again supporting a subduction zone origin for t he LRM meta igneous samples.

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156 Regional correlations between Archean tectonic elements are possible, using the LRM data set to expand on previous geochronology and geochemistry extant for the MHB. Both the Little Belt Mountains and Little Rocky Mountains ar e surface exposures of basement rock within the GFTZ; however, they are geologically very different. The L ittle B elt M ountains appear to have developed in response to subduction between the MHB and the Wyoming craton at ~1.86 Ga, the Little Rocky Mountain s appear to represent earliest Paleoproterozoic to Archean continental crust. The combined LRM MHB dataset also yields more c. 2.4 and 2.8 Ga igneous activity in the MHB, strengthening interpretations of the MHB as an independent entity relative to both t he northerly Hearne craton, and the Wyoming craton to th e south. These published data, in addition to the LRM dataset generated by this study, allows a greater understanding and definition of the MHB.

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157 Table 4 1 Major, trace and rare earth element dat a of orthogneiss samples from the Little Rocky Mountains. Major elements in wt. %, trace and rare earth elements in ppm. Sample MLR 01* MLR 06* MLR 09* MLR 12 MLR 15* MLR 19* LRM 1 LRM 3 LRM 5 LRM 6 LRM 7 LRMGM 08* wt.% Orthogneisses SiO 2 73 74 61 68 61 56 74 63 74 69 72 55 TiO 2 0.04 0.08 0.59 0.59 1.1 1.3 0.21 0.9 0.24 0.47 0.36 1.1 Al 2 O 3 13 14 15 14 14 14 13 15 13 14 13 18 Fe 2 O 3 0.82 0.17 5.2 4.1 6.7 9.0 1.9 6.5 2.1 3.9 3.2 7.4 MnO 0.03 0.01 0.09 0.05 0.09 0.20 0.03 0.09 0.04 0.06 0.04 0.09 MgO 0.23 0.49 3.6 1.3 3.4 4.6 0.38 2.2 0.63 1.3 0.93 4.6 CaO 0.92 1.4 6.1 3.1 4.9 6.4 1.6 4.2 1.6 2.8 1.9 2.5 Na 2 O 3.5 5.1 6.9 3.0 4.0 4.1 3.1 3.5 3.0 3.5 2.9 4.4 K 2 O 7.6 4.9 1.2 4.4 4.6 4.0 4.8 3.5 4.9 3.7 4.9 5.9 P 2 O 5 0.10 0.07 0.21 0 .18 0.42 0.53 0.05 0.33 0.05 0.14 0.11 0.33 LOI 0.46 0.47 0.56 1.1 1.7 1.7 0.52 0.89 0.56 0.71 0.09 0.92 Total 100.46 100.47 100.56 100.13 101.68 101.67 99.11 99.94 100.21 99.64 99.77 100.92 ppm Li N/A N/A N/A 11 6.6 14 4.7 18 6.3 8.1 9.1 11 Sc N/A N/A N/A 7.4 8.7 13 0.03 11 0.12 4.0 4.3 11 Ti N/A N/A N/A 4900 4112 5302 1739 8232 1973 3886 3002 3048 V N/A N/A N/A 54 68 81 20 87 22 50 35 54 Cr N/A N/A N/A 37 49 50 23 60 24 34 37 36 Co N/A N/A N/A 16 11 13 3.2 15 3.5 7.8 6.0 9.3 Ni N/A N/A N/A 19 25 28 8.4 30 8.8 14 15 20 Cu N/A N/A N/A 11 26 29 8.5 32 8.6 9.9 16 13 Zn N/A N/A N/A 49 53 77 30 73 32 48 35 50 Ga N/A N/A N/A 17 16 19 14 20 14 17 17 16 Rb N/A N/A N/A 129 107 112 125 116 128 108 107 98 Sr N/A N/A N/A 212 281 2 52 218 278 233 243 181 179 Y N/A N/A N/A 28 31 50 4.3 39 4.7 13 8.3 24 Zr N/A N/A N/A 24 111 126 41 55 32 65 33 154 Nb N/A N/A N/A 14 16 32 7.0 22 7.8 11 8.3 11 Cs N/A N/A N/A 2.4 2.9 2.7 1.4 1.7 1.1 2.6 2.1 0.80 Ba N/A N/A N/A 1194 1608 868 1735 1449 1901 1369 692 869

PAGE 158

158 Table 4 1. Continued. Sample MLR 01 MLR 06 MLR 09 MLR 12 MLR 15 MLR 19 LRM 1 LRM 3 LRM 5 LRM 6 LRM 7 LRMGM 08 ppm Orthogneisses La N/A N/A N/A 173 109 99 44 114 61 50 59 44 Ce N/A N/A N/A 286 183 202 60 21 0 86 82 94 86 Pr N/A N/A N/A 26 18 22 4.8 22 6.7 7.8 8.2 9.7 Nd N/A N/A N/A 73 57 74 13 74 18 26 25 36 Sm N/A N/A N/A 9.8 9.5 13 1.4 11 1.9 3.9 3.0 7.0 Eu N/A N/A N/A 1.2 1.5 1.7 1.2 1.8 1.5 1.3 0.83 0.98 Gd N/A N/A N/A 7.6 8.0 11 1.3 9.5 1.7 3.3 2.7 6.1 Tb N/A N/A N/A 0.98 1.0 1.5 0.15 1.3 0.19 0.44 0.32 0.85 Dy N/A N/A N/A 5.0 5.6 8.3 0.66 6.9 0.83 2.3 1.5 4.5 Ho N/A N/A N/A 0.95 1.0 1.6 0.14 1.3 0.17 0.44 0.29 0.82 Er N/A N/A N/A 2.7 2.8 4.5 0.51 3.9 0.59 1.3 0.89 2.2 Tm N/A N/A N/A 0 .36 0.40 0.65 0.06 0.55 0.06 0.18 0.11 0.30 Yb N/A N/A N/A 2.1 2.3 4.0 0.42 3.4 0.47 1.1 0.73 1.6 Lu N/A N/A N/A 0.31 0.32 0.56 0.06 0.48 0.07 0.17 0.11 0.23 Hf N/A N/A N/A 0.75 2.8 3.3 1.3 1.6 1.1 1.9 1.0 3.9 Ta N/A N/A N/A 0.85 0.60 1.3 0.31 0.93 0.34 0.43 0.38 0.44 Pb N/A N/A N/A 23 18 19 34 18 37 16 16 15 Th N/A N/A N/A 72 12 15 88 34 127 21 11 4.39 U N/A N/A N/A 0.99 0.73 0.99 0.82 0.85 0.99 0.57 0.55 0.65 *Major elements were analyzed with an incorrect calibration, thus values may be of f by 10%.

PAGE 159

159 Table 4 2 Major, trace and rare earth element data of amphibolite samples from the L ittle R ocky M ountains Major elements in wt. %, trace and rare earth elements in ppm. Sample MLR 03* MLR 07* MLR 10 MLR 11* MLR 14* LRMPG 01* LRMGM 01* LR MGM 02* LRMGM 06* wt. % Amphibolites SiO 2 49 51 53 49 47 50 47 48 50 TiO 2 8.6 8.2 0.90 9.4 7.7 0.97 1.8 1.7 1.2 Al 2 O 3 17 13 12 13 22 11 9.2 8.9 11 Fe 2 O 3 7.2 6.9 12 7.9 7.7 14 17 16 14 MnO 8.5 9.3 0.24 8.4 5.8 0.20 0.24 0.23 0.20 MgO 6.9 6.1 6.8 7.2 6.2 8.1 6.0 6.4 7.9 CaO 1.2 2.0 9.9 1.7 1.2 12 9.0 9.6 11 Na 2 O 0.27 0.42 2.6 0.69 0.63 1.7 1.2 1.6 1.6 K 2 O 0.81 1.0 0.72 0.72 1.7 0.24 0.98 0.32 0.62 P 2 O 5 0.05 0.07 0.08 0.05 0.12 0.10 0.22 0.21 0.16 LOI 0.72 1.4 1.6 2.3 0.87 1.1 0.99 0.69 1.0 Total 99.96 99.86 100.15 100.11 100.03 98.97 93.92 93.83 97.98 ppm Li 2.9 5.4 3.4 5.7 13.6 N/A 7.6 4.3 4.8 Sc 43 30 30 34 43 N/A 42 40 40 Ti 4934 5323 7564 3956 10089 N/A 10715 9523 7017 V 279 222 215 228 421 N/A 422 372 276 Cr 169 172 158 69 69 N/A 62 65 66 Co 49 38 50 39 50 N/A 50 46 47 Ni 70 91 103 70 53 N/A 46 51 45 Cu 125 9 12 50 104 N/A 98 117 65 Zn 99 68 93 98 114 N/A 125 117 96 Ga 15 14 16 14 18 N/A 19 15 17 Rb 15 16 18 29 21 N/A 22 13 18 Sr 96 279 203 177 110 N/A 90 98 164 Y 23 16 18 21 43 N/A 44 39 27 Zr 13 18 9.7 12 26 N/A 22 23 40 Nb 2.5 4.8 4.1 3.5 3.9 N/A 4.2 3.5 6.9 Cs 1.2 2.1 1.5 2.9 1.4 N/A 3.3 1.4 0.92 Ba 95 51 100 103 158 N/A 110 69 178

PAGE 160

160 Table 4 2. Continued. Sample MLR 03 MLR 07 MLR 1 0 MLR 11 MLR 14 LRMPG 01 LRMGM 01 LRMGM 02 LRMGM 06 ppm Amphibolites La 4.8 11 9.2 6.3 7.1 N/A 8.3 5.8 12 Ce 10 21 18 14 17 N/A 20 14 26 Pr 1.4 2.4 2.3 1.9 2.5 N/A 2.9 2.1 3.3 Nd 6.9 9.8 9.5 8.2 12 N/A 14 11 14 Sm 2.2 2.5 2.5 2.5 4.4 N/A 4.4 3.7 3.7 Eu 0.70 0.83 0.80 0.68 1.3 N/A 1.3 1.2 1.1 Gd 2.8 2.6 2.9 2.8 5.5 N/A 5.8 5.0 4.1 Tb 0.52 0.44 0.50 0.51 1.0 N/A 1.1 0.94 0.71 Dy 3.6 2.8 3.0 3.5 7.2 N/A 7.1 6.2 4.4 Ho 0.76 0.56 0.61 0.70 1.5 N/A 1.5 1.3 0.88 Er 2.2 1.5 1. 7 2.0 4.4 N/A 4.4 3.9 2.5 Tm 0.36 0.24 0.25 0.32 0.68 N/A 0.67 0.60 0.39 Yb 2.3 1.4 1.6 2.0 4.4 N/A 4.4 3.9 2.3 Lu 0.35 0.21 0.24 0.28 0.63 N/A 0.65 0.58 0.34 Hf 0.75 0.89 0.72 0.69 1.3 N/A 1.2 1.2 1.3 Ta 0.17 0.35 0.27 0.27 0.27 N/A 0.30 0.25 0. 46 Pb 3.0 4.6 4.0 4.4 6.7 N/A 4.4 2.8 6.1 Th 0.77 3.0 2.6 0.75 1.3 N/A 1.4 1.1 2.6 U 0.15 0.70 0.46 0.75 0.42 N/A 0.46 0.34 0.59 *Major elements were analyzed with an incorrect calibration, thus values may be off by 10%.

PAGE 161

161 Table 4 3 Major, trace a nd rare earth element data of schist samples from the L ittle R ocky M ountains Major elements in wt. %, trace and rare earth elements in ppm. Sample MLR 08* MLR 16 MLR 17 MLR 18 LRMPG 2* LRMPG 3* LRMGM 3* LRMGM 4* LRMGM 5* LRMGM 7* wt. % Schists SiO 2 69 63 61 68 69 66 52 56 56 54 TiO 2 0.28 0.94 1.0 0.63 0.39 0.56 1.6 1.8 1.3 1.2 Al 2 O 3 14 15 15 14 14 15 14 16 14 19 Fe 2 O 3 2.0 6.2 7.3 4.3 1.7 3.3 11 10.0 9.0 7.5 MnO 0.03 0.1 0.1 0.05 0.02 0.05 0.18 0.13 0.12 0.08 MgO 1.2 2.0 2.3 1.4 2.0 2.1 6.3 5.7 5.4 4.1 CaO 3.7 4.0 4.7 2.9 3.8 3.3 7.5 2.3 5.0 4.8 Na 2 O 8.2 3.5 3.5 3.3 7.2 5.3 2.7 3.0 2.9 4.9 K 2 O 0.81 3.5 3.0 4.2 1.5 4.3 4.5 5.3 5.8 4.7 P 2 O 5 0.06 0.35 0.41 0.21 0.11 0.13 0.55 0.54 0.43 0.31 LOI 0.53 0.67 2.2 0.93 0.23 0.52 3.2 2 .4 1.1 1.1 Total 100.52 98.85 100.18 99.99 100.23 100.52 103.18 102.39 101.05 101.15 ppm Li 2.4 18 21 14 6.8 N/A 17 16 17 7.5 Sc 2.1 11 14 6.4 2.9 N/A 14 13 11 6.6 Ti 778 8302 9331 4965 1042 N/A 5751 5609 4754 3181 V 19 87 100 57 21 N/A 102 90 79 49 Cr 7.9 47 61 36 8.1 N/A 57 49 50 17 Co 3.3 30 39 11 4.1 N/A 15 15 12 9.9 Ni 5.0 28 33 18 7.0 N/A 32 29 25 15 Cu 5.1 24 46 16 17 N/A 43 28 22 31 Zn 22 77 88 49 18 N/A 86 64 70 45 Ga 17 20 21 18 16 N/A 19 21 19 17 Rb 11 111 116 12 4 15 N/A 110 112 131 56 Sr 663 270 288 233 501 N/A 302 292 244 295 Y 2.4 44 54 23 5.2 N/A 50 41 41 9.0 Zr 17 48 48 28 57 N/A 60 20 120 75 Nb 1.2 26 25 12 2.2 N/A 24 23 18 6.9 Cs 0.79 1.5 5.2 1.4 0.45 N/A 3.5 3.84 1.91 1.2 Ba 186 1553 1123 1537 263 N/A 1002 1314 1385 848

PAGE 162

162 Table 4 3. Continued. Sample MLR 08 MLR 16 MLR 17 MLR 18 LRMPG 2 LRMPG 3 LRMGM 3 LRMGM 4 LRMGM 5 LRMGM 7 ppm Schists La 7.7 97 344 86 39 N/A 89 137 126 46 Ce 13 202 558 152 62 N/A 181 247 220 76 Pr 1.4 21 50 15 5.8 N/A 20 25 24 7.1 Nd 5.1 70 140 48 18 N/A 72 82 77 23 Sm 0.96 12 19 7.2 2.8 N/A 13 13 12 3.4 Eu 0.55 1.7 2.0 1.4 0.81 N/A 1.6 1.7 1.7 1.0 Gd 0.58 9.6 15 5.8 1.9 N/A 11 11 10 2.9 Tb 0.08 1.4 1.9 0.77 0.20 N/A 1.6 1.4 1.3 0.35 Dy 0.41 7.4 9.5 4.0 0.91 N/A 8.3 7.4 7.0 1.8 Ho 0.07 1.4 1.8 0.77 0.15 N/A 1.6 1.4 1.3 0.33 Er 0.19 4.1 5.0 2.2 0.42 N/A 4.7 4.0 3.7 0.77 Tm 0.02 0.60 0.69 0.30 0.05 N/A 0.65 0.54 0.51 0.13 Yb 0.14 3.7 4.2 1.8 0.28 N/A 3.9 3.3 3.0 0.67 Lu 0.02 0.53 0.59 0.26 0.04 N/A 0.53 0.46 0.42 0.11 Hf 0.54 1.4 1.4 0.78 1.5 N/A 1.7 0.72 3.1 2.0 Ta 0.02 1.6 1.7 0.40 0.08 N/A 0.94 0.86 0.66 0.24 Pb 10 27 36 18 12 N/A 22 18 29 18 Th 1.6 68 125 19 10 N/A 14 53 35 3.8 U 0.11 0.73 0.83 0.46 0.69 N/A 0.64 0.53 0.8 6 0.26 *Major elements were analyzed with an incorrect calibration, thus values may be off by 10%.

PAGE 163

163 Table 4 4 Major, trace and rare earth element data of paragneiss samples from the L ittle R ocky M ountains Major elements in wt. %, trace and rare earth elements in ppm. Sample MLR 04* MLR 05 MLR 13 wt. % Paragneisses SiO 2 62 68 69 TiO 2 1.1 0.64 0.58 Al 2 O 3 13 14 14 Fe 2 O 3 9.1 8.2 4.1 MnO 0.09 0.1 0.05 MgO 7.2 3.0 1.3 CaO 1.8 1.2 2.6 Na 2 O 1.7 1.5 3.3 K 2 O 3.7 2.5 4.3 P 2 O 5 0.16 0.05 0.19 LOI 1.1 0.63 0.80 Total 101.10 100.33 100.48 ppm Li 22 11 10 Sc 14 14 7.1 Ti 3944 3565 4135 V 121 107 61 Cr 134 125 35 Co 20 19 9.8 Ni 61 50 20 Cu 38 40 20 Zn 81 83 51 Ga 17 16 17 Rb 83 82 130 Sr 115 108 220 Y 21 25 34 Zr 94 8 0 176 Nb 9.8 8.7 17 Cs 5.1 1.3 1.5 Ba 435 639 1274 La 38 60 102 Ce 74 113 186 Pr 7.9 12 19 Nd 28 43 63 Sm 5.2 7.9 11 Eu 1.1 1.5 2.0 Gd 4.3 6.6 8.8 Tb 0.60 0.81 1.2 Dy 3.4 4.4 6.3 Ho 0.66 0.84 1.2 Er 1.9 2.5 3.2 Tm 0.30 0.36 0. 44 Yb 1.9 2.4 2.7

PAGE 164

164 Table 4 4. Continued. Sample MLR 04 MLR 05 MLR 13 ppm Paragneisses Lu 0.28 0.35 0.38 Hf 2.7 2.3 4.7 Ta 0.67 0.48 0.62 Pb 15 19 19 Th 13 20 24 U 2.3 2.5 0.96 *Major elements were analyzed with an incorrect calibratio n, thus values may be off by 10%

PAGE 165

165 Table 4 5 Orthogneiss and amphibolite samples L ittle R ocky M ountains zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf i sotope data reported (Ma). Sample Number of Grains 207 Pb/ 206 Pb Error (IA) (IA) (T) (T) (T) Std. Hf Model Age Std. Number Discordance 10% Age Highest Lowest Highest Lowest Average Dev. (DM) a Dev. Orthogneisses MLR 01 20 2420 12 2.7 16.1 2.2 11.6 7.4 2.5 2973 95 MLR 06 31 2416 9 5.5 10.5 5.9 10.0 7.8 1.2 3002 42 MLR 15 43 2776 9 7.7 14.6 7.9 14.5 10.5 1.8 3415 64 MLR 19 56 2793 5 8.7 14.3 8.3 14.4 10.7 1.9 3439 71 LRM 3 13 2800 91 12.1 18.4 11.4 14.2 12.9 0.8 3526 30 LRM 5 8 2827 39 6.7 12.5 6.6 11.7 9.4 1 .4 3416 57 LRM 6 2 N/A N/A 5.8 9.1 N/A N/A N/A N/A 3313 70 LRM 7 12 2769 37 9.2 15.6 8.4 14.0 11.1 1.4 3436 58 MLR 09 24 3199 8 5.2 11.5 3.7 2.3 0.0 1.7 3392 65 Amphibolites MLR 03 32 3006 7 3.6 9.8 4.2 6.3 1.2 2.35 3279 94 MLR 10 5 N/A N/A 2.4 4.4 N/A N/A N/A N/A 3402 82 (IA) Z ircons reduced to individual U Pb age. *zircons reduced to calculated 207 Pb 206 Pb crystallization age. a DM model ages were calculated using the model of Mueller et al. (2008).

PAGE 166

166 Table 4 6 Schist and pa ragneiss samples L ittle R ocky M ountains zircon LA ICP MS U Pb data 207 Pb/ 206 Pb ages and Hf i sotope data reported (Ma). Sample Number of Grains Age Age Age (IA) (IA) Hf T (DM) a Hf T (DM) a Number Discordance 10% Range Population Population Highest Lowest Highest Lowest Schists MLR 08 18 2614 3315 3188 N/A 5.6 15.7 3441 3274 MLR 17 52 2569 2835 2794 2806 6.9 18.5 3685 3318 MLR 18 28 2702 2799 2776 N/A 8.5 16.0 3625 3437 Paragneisses MLR 13 23 2644 2825 2798 2755 7.7 13.3 3531 3315 MLR 04 4 1782 2594 N/A N/A 8.0 16.0 3005 2565 MLR 05 7 2286 3203 2580 3204 1.1 7.1 3437 2749 (IA) Z ircons reduced to i ndividual U Pb age. *zircons reduced to calculated 207 Pb 206 Pb crystallization age. a DM model ages were calculated using the model of Mueller et al. (2008).

PAGE 167

167 Table 4 7 Whole rock LA ICP MS Nd isotope data reported (Ma) Little Rocky Mountains Sample Sm Nd 143 Nd/ 144 Nd (0) (T) Nd Model Age Number (ppm) (ppm) (DM) a Orthogneisses MLR 01 2.31 9.34 0.511596 20.2 5.7 3.551 MLR 06 1.72 10.28 0.510748 36.7 7.3 3.151 MLR 09 2.92 15.89 0.510789 35.9 0.9 3.418 MLR 12 9.77 73. 44 0.510065 50.2 N/A 3.458 MLR 15 9.52 57.23 0.510372 44.0 10.0 3.676 MLR 19 13.33 74.44 0.510478 42.0 10.5 3.801 LRM 1 1.36 13.20 0.509820 54.8 N/A 3.279 LRM 3 11.38 74.01 0.510360 44.3 7.2 3.443 LRM 5 1.86 18.39 0.509797 55.3 6.5 3.28 1 LRM 6 3.88 26.24 0.510292 45.6 N/A 3.430 LRM 7 3.00 24.54 0.510053 50.3 6.8 3.300 LRMPG 03 2.01 11.87 0.510440 42.7 N/A 3.643 LRMGM 05 12.07 77.19 0.510341 44.7 N/A 3.519 LRMGM 08 7.01 35.89 0.510788 35.9 N/A 3.690 Amphibolites MLR 03 2.18 6.91 0.512424 4.0 2.0 N/A MLR 07 2.50 9.83 0.511552 21.0 N/A 3.993 MLR 10 2.49 9.48 0.511716 18.0 N/A 3.861 MLR 11 2.53 8.24 0.512031 11.7 N/A N/A MLR 14 4.36 12.14 0.512783 3.0 N/A N/A LRMPG 01 2.47 7.76 0.512448 3.6 N/A N/A LRMGM 01 4.43 13.89 0.512526 2.0 N/A N/A LRMGM 02 3.74 10.72 0.512825 3.8 N/A N/A LRMGM 06 3.66 13.78 0.512002 12.3 N/A 3.061 Schist MLR 08 0.96 5.06 0.510768 36.3 N/A 3.604 MLR 16 11.83 69.55 0.510424 43.2 N/A 3.679 MLR 17 18.52 140. 44 0.510041 50.7 N/A 3.467 MLR 18 7.25 48.21 0.510266 46.3 N/A 3.507 LRMPG 02 2.79 18.49 0.510235 46.7 N/A 3.557 LRMGM 03 12.91 71.80 0.510540 40.8 N/A 3.720 LRMGM 04 12.96 81.83 0.510341 44.7 N/A 3.558 LRMGM 07 3.38 22.81 0.510262 46.2 N/A 3.475 Paragneiss MLR 04 5.22 28.02 0.511089 30.1 N/A 2.985 MLR 05 7.95 43.07 0.511030 31.2 N/A 3.040 MLR 13 10.68 63.47 0.510393 43.6 N/A 3.690 *Orthogneiss and amphibolite samples reduced to 207 Pb 206 Pb age (Table 4 5). a DM model ag es were calculated using the model of DePaolo (1981).

PAGE 168

168 Figure 4 1 Generalized map of Precambrian basement provinces of southwestern Laurentia (after Ross et al., 1991; Condie, 1992; Doughty et al., 1998; Vogl et al., 2004; Foster et al., 2006, 2012). Exposures of basement in Laramide style uplifts are shown in the dark grey shaded areas.

PAGE 169

169 Figure 4 2. Location of the Little Rocky Mountains relative to a generalized depiction of Cenozoic alkaline rock occurrences in the Montana alkali province (afte r Hearn et al., 1989). The limits of the Great Falls tectonic zone (shown in pink) are not well defined and are based on aeromagnetic data from Sims et al. (2004).

PAGE 170

170 Figure 4 3. Map of the Little Rocky Mountains (after Hearn et al., 1989). Sample loca tions are shown as yellow stars.

PAGE 171

171 Figure 4 4. Photomicrographs of selected meta plutonic samples from the Little Rocky Mountains. Mineral abbreviations are after Whitney and Evans (2010). Plane light (PPL) mircrographs are on the left, and Cross Pol arized light (XPL) micrographs are on the right. Sample MLR 03 (A & B) is an amphibolite, and contains abundant subhedral amphibole grains, anhedral plagioclase, and anhedral, lobate quartz. Sample LRMGM 0 2 (C & D) is similar, but with edral garnet grains. Garnet is rich in quartz and feldspar inclusions. Sample LRMGM 0 8 (E & F) is a biotite granitic orthogneiss, showing biotite with some chlorite alteration, recrystallized quartz, and somewhat altered feldspar. Scale bars represent 1 mm.

PAGE 172

172 Figure 4 5. Photomicrographs of selected meta supracrustal samples from the Little Rocky Mountains. Mineral abbreviations are after Whitney and Evans (2010). PPL mircrographs are on the left, and XPL micrographs are on the right for the first two samples. Sample MLR 13 (A & B) is a staurolite chloritoid quartzofeldspathic paragneiss. Staurolite is poikioblastic, containing numerous inclusions of amphibole, feldspar, and quartz. Chloritoid defines a foliation in the sample, as well as composition al banding between chloritoid rich and poor layers. Titanite occurs as an accessory phase, commonly near staurolite. Sample MLR 04 (C & D) is a garnet biotite quartzofeldspathic augen gneiss. Biotite defines as strong foliation, supported by ribbons of r ecrystallized quartz. Garnet occurs as subhedral grains up to 1mm in diameter, with numerous quartz inclusions. Sample LRMGM 0 7 (E; XPL) and sample LRMPG 0 2 (F; XPL) are fine grained biotite quartzofeldspathic schist. Biotite and a weak compositional va riation define banding. The yellow lines in (E) show the biotite defined foliation. Scale bars represent 1 mm.

PAGE 173

173 Figure 4 6. Schist, orthogneiss and paragneiss samples plotted on an alumina saturation index (Shand, 1943; modified by Frost et al., 2001). Schists (green triangles), orthogneisses (blue diamonds), and paragneisses (red squares).

PAGE 174

174 Figure 4 7. Ternary CaO MgO Al2O3 variation diagram showing the expected fields of meta igneous and meta sedimentary rocks for all schists (circles) and para gn eisses (squares) (after Leyreloup et al., 1977).

PAGE 175

175 Figure 4 8. Primitive for (A) orthogneiss and amphibolite samples and (B) detrital schistose and paragneiss samples. Relative enrichment of large ion lithophile elements (LILs) to high field strength elements (HFSE) suggests formation in an arc environment.

PAGE 176

1 76 Figure 4 9. Chondrite normalized diagram (McDonough and Sun, 1995) for orthogneiss and amphibolite samples using data from Tables 4 1 and 4 2. Orthogneiss samples show elevated concentrations of light versus heavy rare earth elements.

PAGE 177

177 Figure 4 10. Concordia diagrams showing U Pb data showing spread of zircons and weighted mean age for orthogneiss samples MLR 01 (A) and MLR 06 (B). error.

PAGE 178

178 Figure 4 11. Concordia diagrams showing U Pb data for orthogneiss samples MLR 15 (A) and MLR error.

PAGE 179

179 Figure 4 12. ( A) Concordia diagrams plotting composite U Pb data for orthogneiss sample LRM error (B) Expanded view of upper end of concordia showing the oldest, most concordant zircons selected for regress ion.

PAGE 180

180 Figure 4 13. Concordia diagram showing U Pb data for orthogneiss sample LRM 5.

PAGE 181

181 Figure 4 14. Concordia diagram showing U Pb data for orthogneiss sample LRM 7. Each ellipse repre

PAGE 182

182 Figure 4 15. (A) Concordia diagrams plotting composite U Pb data for orthogneiss sample MLR 09. Red dashed lines indicate trends within the analyses. (B) Expanded view of upper end of concordia showing t he oldest, most concordant zircons selected for weighted average. Each ellipse represents a

PAGE 183

183 Figure 4 16. (A) Concordia diagrams plotting composite U Pb data for amphibolite sample MLR 03. Red dashed lines indica te trends within the analyses. (B) Expanded view of upper end of concordia showing the oldest, most concordant zircons selected for weighted average. Each ellipse represents a

PAGE 184

184 Figure 4 17. Probability density pl analyses of the schist and paragneiss samples. Orthogneiss ages are included for comparison. The orange vertical bar represents a time frame when there was both regional metamorphism and igneous crystalli zation occurring in the GFTZ (Chapter s 2 and 3). Blue vertical bars represent crystallization ages from the MHB and GFTZ (Chapter s 2 and 3). Red vertical bars represent crystallization ages from the Wyoming craton. Gray vertical lines represent orthogneis s and amphibolite samples from the LRM.

PAGE 185

185 Figure 4 18. (A) Lu Hf evolution diagram showing the range for the orthogneisss (blue diamond) amphibolites (orange circle) schistose ( green triangle ), and paragneisses ( red square). (B) Sm Nd evolution diagr values for the orthogneisss ( blue diamond ) and the amphibolite (orange Teton Range, BTM Beartooth Mountains, MR Madison Range, TRM Tobacco Root Mountains (ci tations can be found in the text). Zone of lower crustal ratios calculated using Rudnick and Gao (2003) (dotted lines).

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186 Figure 4 19. Published Archean to earliest Proterozoic U Pb ages of zircon from the northern Wyoming craton (citations in the text) compared to published, ages for the MHB (citations in text). Wyoming yellow bars, Medicine Hat purple stars. Crystallization ages of orthogneisss from this study blue stars, from amphibolite orange star, from paragneisses (detrital) pink bars, from schists (detrital) green bars. Prominent age peaks from Table 4 6 black bars.

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187 CHAPTER 5 SUMMARY/ CONCLUSIONS Geochronology and geochemistry of crustal xenoliths and samples from the Little Rocky Mountains (LRM) within the Great Falls tectonic zone (GFTZ) provide new insight into the complex history and heterogeneity of the generally buried Precambrian crust in this region. Xenoliths were entrained in magma by Cenozoic volcanics in the Montana Alkali Province and recovered from two Grassrange diatr emes, which yielded granitoid, schist, and quartzite/sandstone xenoliths; from within the Missouri Breaks Diatremes (Big Slide, Little Sand Creek, and Robinson Ranch localities); from the Bearpaw Mountains at Lloyd Divide, and from the Highwood Mountains. These samples included meta granitoids, orthogneisses, amphibolites, one granofels, one mafic granulite, schistose, and paragneiss xenoliths. In addition, samples of Archean basement rocks exposed within the Little Rocky Mountains were also collected (me ta plutonic orthogneisses and amphibolites as well as meta supracrustal rocks as schists and gneisses). All of the granitoid and meta igneous xenoliths show relative high field strength element (HFSE) depletions and large ion lithophile (LIL) element enric hments characteristic of subduction zone magmatism The limited enrichment of heavy rare earth elements (HREE) relative to primitive mantle values reflects the source mineralogy, likely lower crustal materials where garnet is r esidual after melt extractio n. Zircons from the orthogneiss samples from the Grassrange (Chapter 2) fall into two age populations. The first population is represented by magmatic zircons that yield minimum crystallization ages of ~2.5 Ga. The second age population within the meta plutonic xenoliths is represented by metamorphic zircon growth yielding minimum

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188 metamorphic ages that range between ~1.77 and 1.84 Ga. Two granitoid samples from the Grassrange yield minimum igneous crystallization ages of ~1.73 and 1.74 Ga. Zircon Hf an d whole rock Sm Nd isotopic data for the Archean meta plutonic xenoliths indicate that juvenile and reworked crustal material mixed to varying degrees during crustal formation at ~2.5 Ga. The Hf and Nd isotopic data for the ~1.7 Ga granit oid samples indic ate a mixture of older crust (e.g., the ~2.5 Ga crust in the older xenoliths) and a more juvenile source. This is in contrast to exposures of 1.8 1.9 Ga igneous material in the Little Belt Mountains, which involve far higher proportions of juvenile materi al, hypothesized to represent the arc formed during closure of the Little Belt ocean (Mueller et al., 2002). Similar to data from the Grassrange, granitoid and meta plutonic xenoliths from the Missouri Breaks yield two categories of xenoliths. The first category includes meta plutonic samples with two age populations. The first age population is represented by magmatic zircons that yield minimum igneous crystallization ages that range between ~2.43 and 2.65 Ga, extending the range of Archean ages from wi thin the GFTZ. The second age population is represented by metamorphic zircon growth yielding minimum metamorphic ages that range from 1.72 to 1.75 Ga. The second category of xenoliths include samples that yield a single age population, represented by ma gmatic zircons with calculated minimum igneous crystallization ages of ~1.73 Ga to ~1.89 Ga. Similar to the Grassrange, z ircon Hf and whole rock Sm Nd isotopic data for the Archean meta plutonic xenoliths indicate that juvenile and reworked crustal materi al mixed to varying degrees during crustal formation between ~2. 4 and 2.7 Ga. The only exception to this is a single sample at 2.68 Ga, which seems to primarily be a crustal melt. Unlike the

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189 xenoliths from the Grassrange, t he Hf and Nd isotopic data for the Paleoproterozoic xenoliths from the Missouri Breaks, Bearpaw, and Highwood Mountains are generally similar to Hf and Sm Nd values from the L BM The data reveal xenoliths which, similar to the LBM, involve far higher proportions of juvenile material an d might represent further arc material formed during closure of the Little Belt ocean ( e.g., Mueller et al., 2002). The range of earliest Paleoproterozoic to Archean igneous crystallization? ages from the two xenolith suites are coincident with documented crystallization ages in the M edicine H at B lock (MHB) suggesting that the buried crust sampled by the xenoliths may represent reworked MHB material The metamorphic ages of the Archean xenoliths and the igneous crystallization ages of the Paleoproterozoi c xenoliths correlate with ~1.86 Ga ages from the Little Belt Mountains (LBM) and the tectonic event associated with the collision of the MHB with the Wyoming craton to form the GFTZ. The range of ages in detrital zircons from meta supracrustal xenoliths (schists, paragneisses, etc.) likewise record crustal growth at the times documented in the meta plutonic suites. These data, along with initial Hf isotop ic ratios for Archean detrital zircons, provide no evidence for crust as old as that which characteri zes the northern Wyoming Province (2.8 3.5 Ga ). The crustal xenoliths from this study support the hypothesis that the GFTZ formed due to oceanic plate subduction and arc magmatism that initialized prior to ~1.9 Ga, and continued to 1.78 Ga. The boundary then evolved into a more transpressional continent continent collision by 1.77 Ga as Wyoming moved east towards final collision with the Superior and Hearne provinces as well as the MHB (Dahl et al., 1999; Mueller et al., 2000; Mueller et a l., 2002; Muelle r et al., 2005). There

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190 was likely a period of crustal thickening and metamorphism related to terminal collision followed by crustally derived granitoid magmatism, recorded by the ~1.73 Ga granitoid xenoliths from the Grassrange (Chapter 2). Samples from the Little Rocky Mountains (LRM) reinforce the heterogeneous nature of the GFTZ, as well as the presence of reworked MHB material within it. Magmatic ages from meta plutonic rock collected from the LRM are dominantly Neoarchean (~2.8 Ga) in age, with two early Paleoproterozoic (~2.4 Ga) and two Paleoarchean (~3.0 and 3.2 Ga) ages determined. LRM g eochemical data reveal relative HFSE depletions and LIL enrichments characteristic of subduction zone melts, common to continental crustal material The HREE s how limited enrichment relative to primitive mantle values This reflects the source mineralogy, which is likely lower crustal where garnet is residual after melt extraction. Detrital zircon age spectra from meta supracrustal units in the LRM strongly re cord the c. 2.8 Ga crustal growth event, but lack the older detritus observed in Wyoming Craton detrital zircon suites (e.g. Mueller et al., 1998). The paucity of >3.0 Ga detrital zircon is consistent with a MHB affinity for the LRM sample suite. Evidence of a thermal event that affected the LRM is indicated by the previously determined K Ar and Ar Ar ages at c. 1.7 Ga ( Burwash et al ., 1962) This age coincides with metamorphism in the northwestern W yoming craton (Giletti, 1966; Roberts et al., 2002 Brady et al., 2004) These metamorphic ages record the final continental collision between MHB and WY craton and link the LRM into the larger sample suite studied in this work. Data from this study indicate that the LRM is the only exposed region of MHB crus t. These data provide evidence for a tectonic model in which the MHB at least

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191 partially overrode the Wyoming craton. This significantly improves age control and the isotopic fingerprint on the otherwise concealed MHB, and allows for more robust compariso ns to other Archean elements of the region. Based on data from the xenolith suites and from the LRM, it is possible to extend and refine existing tectonic models for the evolution of the GFTZ. Between 1.89 and 1.79 Ga subduction was ongoing between the MHB and the Wyoming craton with subduction beneath MHB ( Figure 5 1A and B). Arc magmatism was active and granitoid and meta plutonic protoliths to the xenoliths were crystallizing. By 1.78 to 1.77 Ga, the terminal phase of subduction between the MHB and Wyoming province was occurring, leading to crustal thickening, deformation, and wide scale metamorphism as the MHB partially overrode the Wyoming craton ( Figure 5 1C). Evidence can be seen in the Tobacco Root Mountains, southwestern Montana, of Wyoming pr ovince material being buried to depths great enough for anatectic melts to have been created at 1.77 Ga (Mueller et al., 2004; 2005). Similarly, based on the MHB affinity of the Archean meta plutonic samples from this study, the Paleoproterozoic magmatic arc was likely built on MHB crust. Further support is provided by a Deep Probe velocity model (Gorman et al., 2002), which cited an upper mantle floating reflector, interpreted to be a subducted slab, dipping beneath the MHB. This subduction would have l ed to the arc magmatism exposed in the LBM (Mueller et al., 2002). However, the model does not work across the entire GFTZ. If the M edicine H at B lock overrode the Wyoming craton in the central GFTZ, then the ~1.7 Ga granitoids from the Grassrange (Chapte r 2) might be expected to have similarities with the anatectic melts from the southwestern GFTZ, which are melts of Wyoming province material. If the ~1.7 Ga Grassrange granitoids

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192 they are 6 to 10.0). This is evidence that portions of the MHB had to have been at depth as well as Wyoming, leading to partial melting of MHB crustal material as well as Wyoming material ( Figure 5 1D; Mueller et al., 1996). When the oceanic sub duction stopped driving crustal thickening, it is likely that post orogenic collapse and extension began to take place ( Figure 5 1E). This would have led to upwelling of mantle and crustal heating and partially melting lower crustal material, possibly lea ding to the ~1.73 and 1.74 Ga ages seen in xenoliths from the Grassrange (Chapter 2). It is possible that the upwelling mantle could have added mafic magmas to the base of the crust (i.e. high velocity layer; Gorman et al., 2002; Barnhart et al., 2012). Continental collision in the GFTZ ended by ~1.73 Ga, but metamorphic ages within xenolith as young as 1.72 Ga, and xenolith crystallization ages as young as 1.70 (Davis et al., 1995) indicates that there was post collision extensional collapse, and related magmatism and metamorphism within the GFTZ. This could be explained by intraplate orogenisis from the accretion of the Yavapai and Mazatzal terrane which occurred between 1.80 to 1.70 Ga, reactivating tectonic structures within the GFTZ, and leading to f urther metamorphism and partial melting within the area. The formation of the GFTZ appears to have concluded by 1.73 Ga with suturing of the Medicine Hat Block with the Wyoming Province. However, it is important to note that the collision of MHB and Wyomi ng coincided with other circum Wyoming province collisions (Mueller et al., 2011). For example, the Cheyenne belt is proposed to be a Proterozoic suture between the Wyoming province and the Yavapai province. This was believed to have occurred during the Paleoproterozoic Medicine Bow orogeny between

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193 1.66 and 1.80 Ga ( Bickford and Boardman, 1984; Sims and Peterman, 1986; Karlstrom and Bowring, 1988; Premo and Van Schmus, 1989; Van Schmus et al., 1993; Chamberlain, 1998; Selverstone et al., 2000; Hill and Bi ckford, 2001; Hill, 2004 ), clearly overlapping in time with the development of the GFTZ. Further, the Trans Hudson orogen suturing the Wyoming and Superior provinces occurred between 1.71 and 1.77 Ga (Karlstrom and Houston, 1984; Nelson et al., 1993; Reso r et al., 1996; Dahl et al., 1999) Collision b etween the MHB, Wyoming Craton, and Superior Province occurred broadly contemporaneously. This is problematic from a geodynamic perspective, as the high angles of intersection between these convergence zones indicate that they were inherently unstable. T he data presented here support the suggestion of Mueller et al. (2005) that the amalgamation of the Wyoming craton to the Superior craton and the Medicine Hat Craton was in part simultaneous, which places sig nificant limits on plate tectonic relations during this very rapid period of continental growth (e.g., Hoffman, 198 8 ; Mueller et al., 2005).

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194 Figure 5 1. Schematic evolution of the G reat F alls tectonic zone based on geologic constraints and chronolog y. (A) map view: ocean subduction between the Wyoming province (WP) and the M edicine Hat Block (M HB ) Each subsequent panel is a cross section, from modern NW to SE across the MHB to the W P (B) 1.89 1.79 Ga, ocean subduction between the W P and the M HB, arc magmatism in the LBM and xenolith crystallization. (C) 1.78 1.77 Ga, terminal collision, deformation and metamorphism. (D) 1.77 1.75 Ga, crustal thickening as MHB overrides W P crust. Anatexis of W P crust in SW GFTZ (Mueller et al., 2004 200 5 ), melt mixing and xenolith crystallization, contact metamorphism of MHB crust. (E) 1.75 1.73 Ga, post orogenic collapse and extension, upwelling mantle, mafic underplating, partially melting lower crustal rocks, magma crystallization. LBM Little Be lt Mtns.; TRM Tobacco Root Mtns.

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195 APPENDIX A METHODS U Pb A nd Hf Isotopic Analysis O f Zircon Samples were crushed and pulverized using a jaw crusher and disk mill then washed over a Gemini water table to obtain a heavy mineral fraction. The heavy miner al fraction was subjected to magnetic separation using a Franz electromagnetic separator and then concentrated by density in methylene iodide. Zircon grains were selected for analysis using an optical microscope and mounted in a 1 inch epoxy round mount w ith natural zircon standard (FC 1). The mounts were polished and cathodoluminescense (CL) images obtained in order to guide U Pb and Lu Hf analyses. U/Pb analyses were conducted at the University of Florida utilizing a New Wave 213 nm laser to ablate mat erial, which was then analyzed using a Nu Plasma multi collector inductively coupled plasma mass spectrometer (MC ICP MS) following methods summarized by Mueller et al. (2008a). Each U/Pb analysis consists of a 20 second background measurement, 30 seconds of measurement during ablation, and a one minute purging period between analyses. Fractionation corrections are calibrated against the FC 1 natural zircon standard (1098 Ma; Mattinson, 2010). Lu Hf analyses were also conducted with the MC ICP MS with eac h analysis consisting of a 20 second background measurement, one minute of measurement during ablation, and a one minute purging period between analyses. The samples were calibrated against the FC 1 natural zircon standard (Black et al., 2003; Woodhead et al., 2004). Initial Hf isotopic compositions are calculated based on the 207 Pb/ 206 Pb age of the grain and the measured Lu/Hf ratio, unless otherwise stated.

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196 Whole Rock Geochemistry A fraction of each sample was coarsely crushed for whole rock geochemistry Fragments with fresh surfaces were selected and then powdered in a puck mill. Five grams of this aliquot were separated for loss on ignition analysis and then fused for major element analysis by XRF (Geoscience Laboratories). For trace elements, appro ximately 50 mg of powdered material was dissolved in a 6:1 ratio of concentrated hydrofluoric to nitric acid (trace metal grade) After dissolution, the samples were dried and re dissolved in optima (Fisher Scientific) hydrochloric acid. After a second d issolution and drying, the samples were re dissolved in 5% Re Rh spiked nitric acid and analyzed on an ICP MS (Element 2, Thermo Scientific) for trace element abundances. Samples were run against USGS standards AGV 1 and G 2 (Table A 1). All standard val ues lie within error (3 7%) of accepted values when calibrated against AGV 1. After trace elements were measured, a portion of the solution was used to extract Nd for isotopic analysis. Rare earth elements (REE) were separated using cation exchange resin AG50W X12, 200 400 mesh and Nd was separated from the other REE with Ln Resin (Eichrom) prior to isotopic analysis. Sm Nd systematics were evaluated using the Sm/Nd ratios from the trace element analyses, which are accurate to +/ 3% (2 ).

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197 Table A 1. Trace Element data for USGS standards AGV 1 and G 2. Sample AGV 1 AGV 1 AGV 1 G 2 G 2 G 2 Average Standard % Average Standard % ppm n = 7 Deviation Error n = 4 Deviation Error Li 11 1.5 12.7 31 2.1 6.7 Sc 12 0.31 2.6 3.2 0.56 17.5 Ti 7565 1836 24.3 3036 361 11.9 V 120 1.5 1.3 37 2.3 6.3 Cr 11 0.46 4.3 8.5 0.87 10.1 Co 15 0.14 0.9 4.2 0.22 5.1 Ni 16 0.33 2.1 3.9 0.17 4.3 Cu 57 1.8 3.1 11 1.6 14.2 Zn 87 1.9 2.1 81 3.8 4.7 Ga 20 0.22 1.1 23 0.66 2.9 Rb 68 0.79 1.2 165 3.3 2.0 Sr 662 6.0 0.9 463 7.5 1.6 Y 20 0.15 0.8 9.7 0.56 5.7 Zr 227 1.3 0.6 148 32 21.9 Nb 15 0.12 0.8 12 0.30 2.4 Cs 1.2 0.02 1.4 1.3 0.05 3.6 Ba 1196 63 5.3 2020 90 4.5 La 38 1.1 2.7 84 3.9 4.6 Ce 69 1.1 1.6 153 9.5 6.2 Pr 8.2 0.15 1.8 16 0.56 3.6 Nd 31 0.92 3.0 50 1.5 3.1 Sm 5.8 0.12 2.1 6.9 0.17 2.4 Eu 1.6 0.03 1.5 1.4 0.07 5.0 Gd 4.8 0.10 2.1 4.7 0.30 6.5 Tb 0.68 0.01 1.6 0.49 0.02 4.4 Dy 3.6 0.06 1.7 2.1 0.08 4.0 Ho 0.67 0.01 1.0 0.35 0.02 5.1 Er 1.9 0.04 2.0 0.95 0. 06 6.1 Tm 0.27 0.01 3.1 0.11 0.00 4.0 Yb 1.7 0.02 1.5 0.66 0.02 3.7 Lu 0.25 0.00 1.7 0.09 0.00 5.1 Hf 5.1 0.08 1.6 3.6 0.71 20.0 Ta 0.89 0.03 3.1 0.79 0.04 4.5 Pb 37 0.38 1.0 33 1.1 3.3 Th 6.3 0.09 1.4 25 0.77 3.1 U 1.9 0.07 3.6 1.9 0.08 4.4

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198 APPENDIX B SUPPLEMENTARY DATA TABLES Object B 1. External supplemental data tables

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199 LIST OF REFERENCES Annesley, I.R., Madore, C., Krogh, T.E., Anonymous, 1992 U Pb geochronology of mylonitic gneiss and foliated granite from the boundary zone of the Wollaston and Peter Lake domains, Saskatchewan : Program with Abstracts Geological Association of Canada; Mineralogical Association of Canada: Joint Annual Meeting v. 17, p. 3 4. Bailey, R.L., 19 74, Geology and ore deposits of the Alder Gulch area, Little Rocky Mountains, Montana : Unpublished Masters of Science Thesis, Montana State Univ., Bozeman, 81 p. Baird, D.J., Nelson, K.D., Knapp, J.H., Walters, J.J., Brown, L.D., 1996, Crustal structure an d evolution of the Trans Hudson orogen: Results from seismic reflection profiling: Tectonics, v. 15, n. 2, p. 416 426. Barnhart, K.R., Mahan, K.H., Blackburn, T.J., Bowring, S.A., Dudas, F.O., 2012, Deep crustal xenoliths from central Montana, USA: Implica tions for the timing and mechanisms of high velocity lower crust formation : Geosphere, v. 8, n. 6, p. 1408 1428. Bickford, M.E., and Boardman, S.J., 1984, A Proterozoic volcanic plutonic terrane, Gunnison and Salida areas, Colorado: Journal of Geology, v. 92, p. 657 666. Bickford, M.E., Van Schmus, W.R., Collerson, K.D., Macdonald, R., 1987, U Pb zircon geochronology project: New results and interpretation, in Summary of investigations, 1987: Saskatchewan Geological Survey Miscellaneous Report 87 4, p. 76 79. Bickford, M.E., 1988, The formation of continental crust; Part 1, A review of principles; Part 2, An application to the Proterozoic evolution of southern North America: Geological Society of America Bulletin, v. 100, p. 1375 1391. Bickford, M.E., Colle rson, K.D., Lewry, J.F., Van Schmus, W.R., Chiarenzelli, J.R., 1990, Proterozoic collisional tectonism in the Trans Hudson Orogen, Saskatchewan: Geology, v. 18, p. 14 18. Bickford M.E., Hamilton, G.L., Wortman, G.L., Hill, B.M., 2001, Archean rocks in the southern Rottenstone Domain: significance for the evolution of the Trans Hudson Orogen : Canadian Journal of Earth Sciences, v. 38, p. 1017 1025. Black, L.P., Kumo, S.L., Williams, I.S., Mundil, R., Davis, D.W., Korsh, R.J., Foudoulis, C., 2003, The applic ation of SHRIMP to Phanerozoic geochronology; a critical appraisal of four zircon standards: Chemical Geology, v. 200, p. 171 188. Blackburn, T.J., Bowring, S., Burdick, S., van der Hilst, R., Mahan, K., Dudas, F., 2010, U Pb thermochronology: 4 dimensio nal imaging of the North American lithosphere: inSights: The EarthScope Newsletter, spring 2010, 2 p.

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200 Blackburn, T.J., Bowring, S.A., Schoene, B., Mahan, K., Dudas, F., 2011, U Pb thermochronology: Creating a temporal record of lithosphere thermal evolutio n: Contributions to Mineralogy and Petrology, v. 162, p. 479 500 Boerner, D.E., Craven, J.A., Kurtz, R.D., Ross, G.M., Jones, F.W., 1998, The Great Falls Tectonic Zone: Suture or intracontinental shear zone?: Canadian Journal of Earth Sciences, v. 35, n. 2, p. 175 183. Bolhar, R., Kamber, B., Collerson, K., 2007, U Th Pb fractionation in Archaean lower continental crust: Implications for terrestrial Pb isotope systematics: Earth and Planetary Science Letters, v. 254, p. 127 145. Brewer, R.A., Vervoort, J., Lewis, R.S., Gaschnig, R.M., Hart, G., 2008, New constraints on the extent of Paleoproterozoic and Archean basement in the Northwest U.S. cordillera, Eos Trans. AGU, 89(53), Fall Meet. Suppl., Abstract T23C 2066. Buhlmann, A.L., Cavell, P., Burwash, R.A., Creaser, R.A., Luth, R.W., 2000, Minette bodies and cognate mica clinopyroxenite xenoliths from the Milk River area, southern Alberta: records of a complex history of the northernmost part of the Archean Wyoming craton: Canadian Journal of Earth Sciences, v. 37, p. 1629 1650. Burwash, R.A., Baadsgaard, H., Peterman, Z.E., 1962, Precambrian K Ar dates from the western Canada sedimentary basin : Journal of Geophysical Research, v. 67, n. 4, p. 1617 1625. Carlson, R.W. and Irving, A., 1994, Depletion and enric hment history of subcontinental lithospheric mantle: An Os, Sr, Nd and Pb isotopic study of ultramafic xenoliths from the northwestern Wyoming Craton: Earth and Planetary Science Letters, v. 126, p. 457 472. Chamberlain, K.R., 1998, Medicine Bow Orogeny; t iming of deformation and model of crustal structure produced during continent arc collision, ca. 1.78 Ga, southeastern Wyoming: Rocky Mountain Geology v. 33, n. 2 p. 259 277. Chappell, B.W. and White, J.R., 2001, Two contrasting granite types: 25 years la ter: Australian Journal of Earth Science, v. 48, p. 489 499. Chauvel, C. and Blichert Toft, J., 2001, A hafnium isotope and trace element perspective on menting of the depleted mantle: Earth and Planetary Science Letters, v. 190, p. 137 151. Collerson, K., Hearn, B.C., MacDonald, R.A., Upton, B.G.J., Harmon, R.S., 1989, Composition and evolution of lower continental crust: Evidence from xenoliths in Eocene lavas from the Bearpaw Mountains, Montana, in Continental magmatism: Abstracts of the IAVCEI: New Mexi co Bureau of Mines and Resources Bulletin 131, p. 57.

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201 Condie, K.C., 1992, Proterozoic terranes and continental accretion in southwestern North America: In Condie, K. C., ed. Proterozoic crustal evolution. Amsterdam, Elsevier, chap. 12, p. 447 480. Cox, D. Frost, C., Chamberlain, K., 2000, 2.01 Ga Kennedy dike swarm, southeastern Wyoming: Record of a rifted margin along the southern Wyoming Province: Rocky Mountain Geology, v. 35, n. 1, p. 7 30. Dahl, P.S., Holm, D.K, Gardner, E.T., Hubacher, F.A., Foland, K.A., 1999, New constraints on the timing of Early Proterozoic tectonism in the Black Hills (South Dakota), with implications for docking of the Wyoming province with Laurentia: GSA Bulletin, v. 111, n. 9, p. 1335 1349. Dahl, P.S., Hamilton, M.A., Stern, R.A., Frei, R., Berg, R.B., 2000, In situ SHRIMP investigation of an early Proterozoic metapelite, with implications for Pb Pb step leach dating of garnet and staurolite: Abstracts with Programs Geological Society of America, v. 32, n. 7, p. 297. Davidso n, A., 2008 Late Paleoproterozoic to mid Neoproterozoic history of northern Laurentia; an overview of central Rodinia : Precambrian Research, v. 160, p. 5 22. Davis, W.J., Berman, R., Kyarsgaard, B., 1995, U Pb geochronology and isotopic studies of crustal xenoliths from the Archean Medicine Hat block, northern Montana and southern Alberta: Paleoproterozoic reworking of Archean crust. In 1995 Alberta Basement Transects Workshop. Edited by G.M. Ross. Lithoprobe secretariat, The University of British Columbia Lithoprobe Report 47, p. 330 335. Deer, W.A., Howe, R.A., Zussman, J., 1992, The rock forming minerals, second edition. Pearson Education, Ltd, Essex. 696 p. DePaolo, D.J., 1981, Trace element and isotopic effects of combined wallrock assimilation and fractional crystallization: Earth and Planetary Science Letters, v. 53, p. 189 202. Dhuime, B., Hawkesworth, C., Cawood, P., 2011, When continents formed: Science, v. 331, p. 154 155. Doughty, P.T., Price, R.A., Parrish, R.R., 1998, Geology of Archean base ment and Proterozoic cover in the Priest River complex, northwestern United States, and their implications for Cordilleran structure and Precambrian continent reconstructions: Canadian Journal of Earth Sciences, v. 35, p. 39 54. Downes, H., MacDonald, R., Upton, B., Cox, K., Bodinier, J., Mason, P., James, D., Hill, P., Hearn, B.C., 2004, Ultramafic xenoliths from the Bearpaw Mountains, Montana, USA: Evidence for multiple metasomatic events in the lithospheric mantle beneath the Wyoming craton: Journal of P etrology, v. 45, p. 1631 1662.

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202 Dyson, J.L., 1939, Ruby Gulch mining district, Little Rocky Mountains, Montana. Economic Geology, v. 34, p. 201 213. Eaton, D.W., Ross, G.M., Clowes, R.M., 1999, Seismic reflection and potential field studies of the Vulcan s tructure, Western Canada; a Paleoproterozoic Pyrenees?: Journal of Geophysical Research, v. 104, p. 23,255 23,269. Facer, J., Downes, H., Beard, A., 2009, In situ serpentinization and hydrous fluid metasomatism in spinel dunite xenoliths from the Bearpaw Mountains, Montana, USA: Journal of Petrology, v. 50, p. 1443 1475. Faure, G. and Mensing, T.M., 2005, Isotopes Principles and Applications, Third edition. John Wiley & Sons, Inc. Hoboken. Foster, D.A., Mueller, P.A., Mogk, D.W., Wooden, J.L., Vogl, J.J. 2006, Proterozoic evolution of the western margin of the Wyoming craton: implications for the tectonic and magmatic evolution of the northern Rocky Mountains: Canadian Journal of Earth Sciences, v. 43, n, p. 1601 1619. Foster, D.A., Mueller, P.A., Heathe rington, A. Kalakay, T.J., 2012, Lu Hf systematics of magmatic zircons reveal a Proterozoic crustal boundary under the Cretaceous Pioneer batholith, Montana: Lithos, v. 142 143, p. 216 225. Frost, C.D., 1993, Nd isotopic evidence for the antiquity of the W yoming province: Geology, v. 21, p. 351 354. Frost, C.D., Frost, B.R., Chamberlain, K.R., Hulsebosch, T.P., 1998, The Late Archean history of the Wyoming province as recorded by granitic magmatism in the Wind River Range, Wyoming: Precambrian Research, v. 89, p. 145 173. Frost, B.R., Barnes, C.G., Collins, W.J., Arculus, R.J., Ellis, D.J., Frost, C.D., 2001, A geochemical classification for granitic rocks: Journal of Petrology, v. 42, p. 2033 2048. Frost, C.D. and Fanning, C.M., 2006 a Archean geochronolog ical framework of the Bighorn Mountains, Wyoming : Canadian Journal of Earth Sciences, v. 43, n. 10, p. 1399 1418. Frost, C.D., Frost, B.R., Kirkwood, R., Chamberlain, K., 2006b, The tonalite trondhjemite granodiorite (TTG) to granodiorite granite (GG) tran sition in the late Archean plutonic rocks of the central Wyoming Province: Canadian Journal of Earth Sciences, v. 43, p. 1419 1444. Giletti, B., 1966, Isotopic ages from southwestern Montana: Journal of Geophysical Research, v. 71, 4029 4036.

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203 Gordon, T.M. Hunt, P.A., Bailes, A.H., Syme, E.C., 1990, U Pb ages from the Flin Flon and Kisseynew belts, Manitoba; chronology of crust formation at an early Proterozoic accretionary margin: Special Paper, Geological Association of Canada, v. 37, p. 177 199. Gorman A, Clowes, R, Ellis, R., Henstock, T., Spence, G., Keller, G., 2002, Deep Probe: imaging the roots of western North America: Canadian Journal of Earth Sciences v. 39, n. 3, p. 375 398. ., Xu, X., Zhou, X., 2002, Zircon chemistry and magma mixing, SE China: In situ analysis of Hf isotopes, Tonglu and Pingtan igneous complexes: Lithos, v. 61, p. 237 269. Guevara, V.E., Baldwin, J.A., Foster, D.A., Lewis, R.S., 2012 From peak metamorphism to orogenic collapse: insights into the exhumation history of the Clearwater metamorphic core complex : Mineralogical Magazine v. 76, p. 1788. Harms, T., Brady, J., Burger, H., Cheney, J., 2004, Advances in the geology of the Tobacco Root Mountains, Mont ana, and their implications for the history of the northern Wyoming Province, in Brady, J.B., et al., eds., Precambrian geology of the Tobacco Root Mountains, Montana: Geological Society of America Special Paper 377, p. 227 243. Hearn, B.C., Jr., Marvin, R .F., Zartman, R.E., Naeser, C.W., 1977, Geochronology of igneous activity in the north central alkalic province. Geological Society of America Abstracts with Programs, v. 9, p. 732. Hearn, B.C., Jr Dudas, F.O., Eggler, D.H., Hyndman, D.W., O'Brien, H.E., 1989, Volcanism and plutonism of western North America; Volume 2, Montana high potassium igneous province, Washington, DC, United States (USA): American. Geophysical. Union. Heimlich, R.A. and Banks, P.O., 1968, Radiometric age determinations, Bighorn Mou ntains, Wyoming: American Journal of Science v. 266, n. 3, p. 180 192. Henstock, T.J., Levander, A., Snelson, C.M., Keller, G.R., Miller, K.C., Harder, S.H., Gorman, A.R., Clowes, R.M., Burianyk, M.J.A., Humphreys, E.D., 1998, Probing the Archean and Prote rozoic Lithosphere of Western North America: GSA Today, v. 8, n. 7, p. 1 5. Hill, B.M., and Bickford, M.E., 2001, Paleo Proterozoic rocks of central Colorado: Accreted arcs or extended older crust?: Geology, v. 29, p. 1015 1018, Hill, B.M., 2004, Paleoprot erozoic of central Colorado: Island arcs or rifted older crust? [Ph.D. thesis]: Syracuse, New York, Syracuse University, 145 p.

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204 Holm, D. and Schneider, D., 2002, 40 Ar/ 39 Ar evidence for ca. 1800 Ma tectothermal activity along the Great Falls tectonic zone, central Montana: Canadian Journal of Earth Sciences, v. 39, p. 1719 1728. Hoffman, P.F., 1988, United plates of America, the birth of a craton: Early Proterozoic assembly and growth of Laurentia: Annual Review of Earth and Planetary Sciences, v. 16, p. 543 603. Hoffman, P.F., 1989, Speculations on Laurentia's first gigayear (2.0 to 1.0 Ga): Geology, v. 17, p. 135 138. Hoffman, P., 1990 Archean continental plat es; old and young mantle roots: Nature v. 347, p. 19 20. Jansen, A.C., Vervoort, J.D., Reed, S ., Anonymous, 2011 Precambrian basement rocks of the Clearwater metamorphic complex: a new piercing point along the western margin of Laurentia : Abstracts with Programs Geological Society of America, v. 43, p. 491. Joswiak, D., 1992, Composition and evolu tion of the lower crust, central Montana; evidence for granulite xenoliths [M.S. thesis]: Seattle, University of Washington, 154 p. Karlstrom, K.E. and Houston, R.S., 1984, The Cheyenne Belt: Analysis of a Proterozoic suture in southern Wyoming: Precambria n Research, v. 25, p. 415 446. Karlstrom, K.E. and Bowring, S.A., 1988, Early Proterozoic orogeny assembly of tectonostratigraphic terranes in southwestern North America: Journal of Geology, v. 96, p. 561 576. Kirkwood, R., 2000, Geology, geochronology and economic potential of the Archean rocks in the western Owl Creek Mountains, Wyoming : Dissertation at University of Wyoming, Laramie, WY, United States. Kleinkopf, M.D., Witkind, I.J., Keefer, W.R., 1972, Aeromagnetic, Bouguer gravity, and generalized geol ogic maps of the central part of the Little Belt Mountains, Montana: Geophysical Investigations Map: 4. U. S. Geological Survey. Knechtel, M., 1959, Stratigraphy of the Little Rocky Mountains and encircling foothills, Montana. U.S. Geological Survey Bullet in 1072 N, p. 723 752. Krogh, T., Kamo, S.L., Hanley, T.B., Hess, D.F., Dahl, P.S., et al., 2011, Geochronology and geochemistry of Precambrian gneisses, metabasites, and pegmatite from the Tobacco Root Mountains, northwestern Wyoming craton, Montana: Cana dian Journal of Earth Sciences, v. 48. n. 2, p. 161 185.

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205 Lemieux, S., Ross, G.M., Cook, F.A., 2000, Crustal geometry and tectonic evolution of the Archean crystalline basement beneath the southern Alberta Plains, from new seismic reflection and potential field studies: Canadian Journal of Earth Sciences, v. 37, n. 11, p. 1473 1491. Lewry, J.F., Thomas, D.J., MacDonald, R., Chiarenzelli, J., 1987, Structural relations in accreted terranes of the Trans Hudson Orogen, Saskatchewan: telescoping in a collisiona l regime?: Geological Association of Canada meeting, Abstracts, v.12, p. 67. Lewry, J.F., Hajnal, Z., Green, A., Lucas, S.B., White, D., Stauffer, M., Ashton, K., Weber, W., Clowes, R., 1994, Structure of the Paleoproterozoic continent continent collision zone: a LITHOPROBE seismic reflection profile across the Trans Hudson Orogen, Canada: In Proceedings of the 5 th International conference on seismic reflection probing of continents and their margins. Edited by R.M. Clowes and A.G. Green. Tectonophysics, v. 232 p. 143 160. Leyreloup, A., Dupuy, C., Andriambololona, R., 1977, Catazonal xenoliths in French Neogene volcanic rocks: Constitution of the lower crust: Contributions to Mineralogy and Petrology, v. 62, p. 283 300. Lucas, S.B., Stern, R.A., Syme, E.C. Reilly, B.A., Thomas, D.J., 1996, Intraoceanic tectonics and development of continental crust: 1.92 1.84 Ga evolution of the Flin Flon Belt, Canada: Geological Society of America Bulletin, v. 108, p. 602 629. MacDonald, R., Upton, B., Collerson, K., Hear n, B.C., James, D., 1992, Potassic mafic lavas of the Bearpaw Mountains, Montana: Mineralogy, chemistry, and origin: Journal of Petrology, v. 33, p. 305 346. Machado, N., 1990, Timing of collisional events in the Trans Hudson Orogen: Evidence from U Pb geo chronology for New Quebec Orogen, the Thompson Belt, and the Reindeer Zone Manitoba and Saskatchewan: In The Early Proterozoic Trans Hudson Orogen of North America Edited by J.F. Lewry and M.R. Stauffer. Geological Association of Canada, Special Paper, v. 37, p. 433 441. Machado, N., Zwanzig, H., Parent, M., 1999, U Pb ages of plutonism, sedimentation, and metamorphism of the Paleproterozoic Kisseynew metasedimentary belt, Trans Hudson Orogen (Manitoba, Canada): Canadian Journal of Earth Science, v. 36, p. 1829 1842. Maniar, P.D. and Piccoli, P.M., 1989, Tectonic discrimination of granitoids: Geological Society of America Bulletin, v. 101, p. 635 643. Marvin, R.F., Witkind, I.J., Keefer, W.R., Mehnert, H.H., 1973, Radiometric a ges of i ntrusive r ocks in the L ittle Belt Mountains, Montana: GSA Bulletin, v. 84, n. 6, p. 1977 1986.

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206 Marvin, R.F., Hearn, B.C., Mehnert, H.H., Naeser, C.W., Zartman, R.E., Lindsey, D.A., 1980, Late Cretaceous Paleocene Eocene igneous activity in north central Montana: Isochron West, v 29, n. 3, p. 5 25. Mattingon, J., 2010, Analysis of the relative decay constants of 235 U and 238 U by multi step CA TIMS measurements of closed system natural zircon samples: Chemical Geology v. 275, p. 186 198. McDonough, W.F. and Sun, S., 1995, The co mposition of the Earth: Chemical Geology, v. 120, p. 223 253. Miller S.H., Hildebrandt, P.K., Erlsev, E.A., Reed, J.C. Jr., 1986 M etamorphic and deformation history of the gneiss complex in the northern Teton Range, Wyoming. In Geology of the Beartooth uplift and adjacent ranges, Montana Geologic Society and Yellowstone Beartooth Research Association Joint Field Conference and Symposium, Red Lodge, Mont., 31 August 1 September 1986, p. 91 110. Mirnejad., H. and Bell, K., 2006, Origin and source evoluti on of the Leucite Hills lamproites: evidence from Sr Nd Pb O isotopic compositions: Journal of Petrology, v. 47, p. 2,463 2,489. Mirnejad., H. and Bell, K., 2008, Geochemistry of crustal xenoliths from the Hatcher Mesa lamproite, Wyoming, USA: Insights i nto the composition of the deep crust and upper mantle beneath the Wyoming craton. Can. Mineral., v. 46, p. 583 596. Mogk, D.W., Mueller, P.A., Weyand, E., Heatherington, A.L., 1990, Middle Archean gneisses of the northern Madison Range, SW Montana: Crus tal genesis in a magmatic arc?: Eos (American Geophysical Union Transactions), v. 17, p. 655. Mogk, D.W., Mueller, P.A., Wooden, J.L., 2004, Tectonic implications of late Archean early Proterozoic supracrustal rocks in the Gravelly Range, SW Montana : Abstr acts with Programs, Geological Society of America, v. 36, p. 507. Mueller, P.A. and Wooden, J.L., 1988, Evidence for Archean subduction and crustal recycling, Wyoming province: Geology, v. 16, p. 871 874. Mueller, P.A., Shuster, R., Wooden, J., Erslev, E., Bowes, D., 1993, Age and composition of Archean crystalline rocks from the southern Madison Range: Implications for crustal evolution in the Wyoming craton: Geological Society of America Bulletin, v. 105, p. 437 446. Mueller, P.M., Heatherington, A., Weya nd, E, Wooden, J., Mogk, D.A.,1995, Geochemical evolution of Archean crust in the northern Madison Range, Montana, evidence from U Pb and Sm Nd systematics: Abstracts with Programs Geological Society of America, v. 27, n. 4, p. 49.

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207 Mueller, P.A., Wooden, J.L., Mogk, D.W., Nutman, A.P., Williams, I.S., 1996, Extended history of a 3.5 Ga trondhjemitic gneiss, Wyoming Province, USA: Evidence from U Pb systematics in zircon: Precambrian Research, v. 78, p. 41 52. Mueller, P.A., Heatherington, A., Kelly, D., Wooden, J.L., Mogk, D., 2000, The Great Falls tectonic zone and its role in the Paleoproterozoic assembly of southern Laurentia: Abstracts with Programs Geological Society of America v. 32, n. 7, p. 318. Mueller, P.A., Heatherington, A.L., Kelly, D.M., W ooden, J.L., Mogk, D.W., 2002, Paleoproterozoic crust within the Great Falls tectonic zone: Implications for the assembly of southern Laurentia: Geology, v. 30, n. 2, p. 127 130. y, K., 2004, Age and evolution of the Precambrian crust of the Tobacco Root Mountains, Montana, in Brady, J.B., et al., eds., Precambrian geology of the Tobacco Root Mountains, Montana: Geological Society of America Special Paper 377, p. 181 202. Mueller, P., Burger, H., Wooden, J., Brady, J., Cheney, J., Harms, T., Heatherington, A., and Mogk, D., 2005, Paleoproterozoic metamorphism in the northern Wyoming Province: Implications for the assembly of Laurentia: Journal of Geology, v. 113, p. 169 179. Mueller P.A. and Frost, C.D., 2006, The Wyoming province: a distinctive Archean craton in Laurentian North America: Canadian Journal of Earth Sciences, v. 43, p. 1391 1397. Mueller, P.A., Kamenov, G.D., Heatherington, A.L., Richards, J., 2008a, Crustal evolution in the Southern Appalachian orogen: Evidence from Hf isotopes in detrital zircons: Journal of Geology, v. 116, n. 4, p. 414 422. Mueller, P.A., Foster, D.A., Gifford, J.N., Wooden, J., Mogk, D.W., 2008b, Tectonic and stratigraphic implications of detrital zircon suites in Cambrian and Precambrian sandstones from the Eastern margin of the belt basin: Northwest Geology, v. 37, p. 61 68. Mueller, P., Wooden, J.L., Mogk, D., Henry, D., Bowes, D.R., 2010, Rapid growth of an Archean continent by arc magmatism: P recambrian Research, v. 183, p. 70 88. Mueller, P.A., Foster, D.A., Gifford, J.N., Vogl, J.J., Wooden, J., Mogk, D.W., 2011, Closure of a Paleoproterozoic ocean recorded by Hf isotopes in zircons: Geological Society of America Abstracts with Programs v. 43 n. 5, p. 323. Mueller, P.A. and Wooden, J.L, 2012, Trace element and Lu Hf systematics in Hadean Archean detrital zircons: Implications for crustal evolution: The Journal of Geology, v. 120, n. 1, p. 15 29.

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208 Nelson, K.D., Baird, D.J., Walters, J.J., Hauck M., Brown, L.D., Oliver, J.E., Ahern, J.L., Hajnal, Z., Jones, A.G., Sloss, L.L., 1993, Trans Hudson orogen and Williston basin in Montana and North Dakota: New COCORP deep profiling results: Geology, v. 21, p. 447 450. haracter and regional significance of Great Falls tectonic zone, East Central Idaho and West Central Montana: American Association of Petroleum Geologists Bulletin. v.69, p. 437 447. Pearce, J.A., 1983. The role of the sub continental lithosphere in magma genesis at destructive plate boundaries: In Thorpe, R.S. (ed.) Continental Basalts and Mantle Xenoliths, Shiva, Nantwich, p. 230 249. Pearce, J.A., Harris, N.B.W., Tindle, A.G., 1984, Trace element discrimination diagrams for the tectonic interpretation o f granitic rocks: Journal of Petrology. v. 25, p. 956 983. Peterman, Z.E., 1981, Archean Gneisses in the Little Rocky Mountains, Montana. Geological Survey Professional Paper 1199 A, p. 1 47. Pilkington, M., Grieve, R.A.F., Rupert, J.D., Halpenny, J.F., 19 92, Gravity anomaly map with shaded relief of gradient of North America. Geologic Survey of Canada, Map 1807A, scale 1:10,000,000. Pirsson, L., 1900, Petrography of the igneous rocks on the Little Belt Mountains: U.S. Geological Survey, 20 th Annual report, Part III, p. 116 134. Premo, W. R. and Van Schmus, W.R., 1989, Zircon geochronology of Precambrian rocks in southeastern Wyoming and northern Colorado: Proterozoic geology of the southern Rocky Mountains: Geological Society of America Special Paper, v. 23 5, p. 13 32. Ray G.E. and Wanless, R.K., 1980 The age and geological history of the Wollaston, Peter Lake, and Rottenstone do mains in northern Saskatchewan: Can adian Journal of Earth Sciences, v. 17, p. 333 347. Rayner, N.M ., Stern, R.A., Bickford, M.E., 2005, Tectonic implications of new SHRIMP and TIMS U Pb geochronology of rocks from the Sask Craton, Peter Lake Domain, and Hearne margin, Trans Hudson Orogen, Saskatchewan : Canadian Journal of Earth Sciences, v. 42, p. 635 657. Resor, P.G., Snoke, A.W., Chamberlain, K.R., 1996, Development of a shear zone bounded block uplift within the middle crust of the Archean Wyoming Province during Proterozoic accretion, Laramie Mts., WY: Abstracts with Programs, Geological Society of America, v. 28, p. 496.

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209 Rober ts, H., Dahl, P., Kelley, S., and Frei, R., 2002, New 207 Pb 206 Pb and 40 Ar 39 Ar ages from SW Montana, USA: Constraints on the Proterozoic and Archean tectonic and depositional history of the Wyoming Province: Precambrian Research, v. 117, p. 119 143. Roger s, G.R. and Hawkesworth, C.J., 1989, A geochemical traverse across the north Chilean Andes: evidence for crust generation from the mantle wedge: Earth and Planetary Science Letters, v. 91, p. 271 285. Ross, G.M., 1991, Precambrian basement in the Canadian Cordillera: an introduction: Canadian Journal of Earth Sciences, v. 28, p. 1133 1139. Ross, G. M. 2002, Evolution of Precambrian continental lithosphere in Western Canada: Results from Lithoprobe studies in Alberta and beyond: Canadian Journal of Earth Sci ences, v. 39, p. 413 437 Rubatto D ., 2002 Zircon trace element geochemistry: partitioning with garnet and the link between U Pb ages and metamorphism : Chemical Geology v. 184 p. 123 138 Rudnick, R.L. and Gao, S., 2003, Composition of the continental c rust: Treatise on geochemistry, v. 3, p. 1 64. Russell, C.W., 1984, Geology of the central portion of the Little Rocky Mountains, Phillips County, Montana: Unpublished Masters of Science Thesis, Univ. Idaho, Moscow, 92 p. Schafer, P.A., 1935 Geology and ore deposits of the Neihart mining district, Cascade County, Montana : Memoir State of Montan a, Bureau of Mines and Geology, v. 62. Sears, J.W. and Link, P.K., 2007, A field guide to the quartzite of Argenta; Neihart Quartzite equivalent?: Northwest Geolo gy, v. 36, p. 221 230. Selverstone, J., Condie, K.C., Van Schmus, W.R., 2000, The crust of the Colorado Plateau; evidence from the xenolithic record: Abstracts with Programs, Geological Society of America, v. 32, p. 386. Shand, S.J., 1943, The Eruptive Ro cks, 2nd edition: New York, John Wiley, 444 p. Shaw, D.M., 1972, Geochemistry in Canada: Earth Science Reviews. v. 8, p. 129 131. Sims, P.K. and Peterman, Z.E., 1986, Early Proterozoic Central Plains orogen: a major buried structure in the north central Un ited States: Geology, v. 14, p. 488 491. Precambrian basement geologic map of Montana; an interpretation of aeromagnetic anomalies. Scientific Investigations Map, United States Geologic Survey.

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210 Taylor, S.R. and McLennan, S.M., 1995, The geochemical evolution of the continental crust: Reviews of Geophysics, v. 33, n. 2, p. 241 265. Thomas, M.D., Sharpton, V.L., Grieve, R.A.F., 1987, Gravity patterns and Precambrian structure in the North American Central Plains: Geology, v. 15, p. 489 492. Thompson, R.N., Morrison, M., Hendry, G., Parry, S., 1984, An assessment of the relative roles of crust and mantle in magma genesis: an elemental approach. Philisophical Transactions of the Royal Society of Lond on. v. A310, p. 549 590. Van Schmus, W.R, Bickford, M.E., Condie, K.C., 1993, Early Proterozoic transcontinental orogenic belts in the United States: Abstracts with Programs, Geological Society of America, v. 25, p. 44. Villaseca, C., Orejana, D., Belous ova, E.A., 2012, Recycled metaigneous crustal sources for S and I type Variscan granitoids from the Spanish Central System batholith: Constraints from Hf isotope zircon composition: Lithos, v. 153, p. 84 93. Villeneuve, M.E., Ross, G.M., Parrish, R.R., Th eriault, R.J., Miles, W., Broome, J., 1993, Geophysical subdivision, U Pb geochronology and Sm Nd isotope geochemistry of the crystalline basement of the Western Canada sedimentary basin, Alberta and northeastern British Columbia: Geological Survey of Cana da Bulletin 447, 86 p. Vogl, J.J., Mueller, P.A., Foster, D.A., Wooden, J.L., 2003, Paleoproterozoic history of the Great Falls tectonic zone; results from an integrated study of basement exposures in the Little Belt Mountains, Montana: Abstracts with Prog rams Geological Society of America, v. 35, n. 6, p. 595 596. Vogl, J., Foster, D., Mueller, P., Wooden, J., Mogk, D., 2004, Lithology and age of pre Belt Precambrian basement in the Little Belt Mountains, Montana: Implications for the role of the Great Falls tectonic zone in the Paleoproterozoic assembly of North America: Northwest Geology, v. 33, p. 15 32. Weed, W.H., 1900, Geology of the Little Belt Mountains, Montana: Twentieth Annual Report of the United States Geological Survey. Part III. Precious M etal Mining Districts. Government Printing Office, Washington, D. C. Weed, W.H. and Pirsson, L.V., 1896, The Geology of the Little Rocky Mountains, Montana. Journal of Geology, p. 399 428. Weiss, R., Vogl, J., Mueller, P., Foster, D., Wooden, J., 2009, Lu Hf analyses of zircons in the Little Belt Mountains suggest Paleoproterozoic subduction in the Great Falls tectonic zone: AGU Fall Meeting, V43E 2321.

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211 Werner, C. D., 1987, Saxonian granulites: a contribution to the geochemical diagnosis of original rocks in high metamorphic complexes: Gerlands Beitrage zur Geophysik, v. 96, p. 271 290. Whitehouse, M.J., Stacey, J.S., Miller, F.K., 1992, Age and nature of the basement in northeastern Washington and northern Idaho: isotopic evidence from Mesozoic and Cenozo ic granitoids. Journal of Geology, v. 100, p. 691 701. Wilson, M.R. and Kyser, T.K., 1988, Geochemistry of porphyty hosted Au Ag deposits in the Little Rocky Mountains, Montana. Economic Geology, v. 83, p. 1329 1346. Wooden, J.L., and Mueller, P.A., 1988 Pb, Sr, and Nd isotopic compositions of a suite of Late Archean, igneous rocks, eastern Beartooth Mountains: Implications for crust mantle evolution: Earth and Planetary Science Letters, v. 87, p. 59 72. Woodhead, J., Hergt, J., Shelley, M., Eggins, S., Kemp, R., 2004, Zircon Hf isotope analysis with an excimer laser, depth profiling, ablation of complex geometries, and concomitant age estimation: Chemical Geology v. 209, n. 1 2, p. 121 135.

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212 BIOGRAPHICAL SKETCH Jennifer was born in Hartford, C onnecticu t in 1983 to Michael and Terry Gifford She is the youngest of three sisters. She graduated from R egional H ebron A ndover M arlborough (RHAM) high school (Hebron, C onnecticut ) in 2001. She received a Bachelor of Science degree in geology from Syracuse Un iversity (Syracuse, N ew Y ork ) in 2005. While attending graduate school at the University of Florida (Gainesville, F lorida ), Jennifer served as a teaching assistant for several courses in the Department of Geological Sciences and as a research assistant fo r Dr. David A. Foster and Dr. Paul May of 2008. She received a doctoral degree in geology from the University of Florida in May of 2013.