Copyright by James V. Jones III 2005

The Dissertation Committee for James V. Jones III Certifies that this is the approved version of the following dissertation:

Proterozoic tectonic evolution of southern Laurentia: New constraints from field studies and geochronology in southern Colorado and northern New Mexico, U.S.A.

Committee:

James Connelly, Supervisor

Mark Cloos

Sharon Mosher

Karl Karlstrom

Kent Condie

Proterozoic tectonic evolution of southern Laurentia: New constraints from field studies and geochronology in southern Colorado and northern New Mexico, U.S.A.

by

James V. Jones III, B.S.; M.S.

Dissertation Presented to the Faculty of the Graduate School of The University of Texas at Austin in Partial Fulfillment of the Requirements for the Degree of

Doctor of Philosophy

The University of Texas at Austin August, 2005 Acknowledgements

I would first like to thank my advisor, Jim Connelly, for five years of support, encouragement, and guidance. I would also like to thank my dissertation committee members Mark Cloos, Sharon Mosher, Karl Karlstrom, and Kent Condie for their time and efforts in providing insightful comments and lively discussion about this work. An extensive group of colleagues and collaborators was always willing to lend time, expertise, and information and includes (among others and in no particular order) Karlstrom, Mike Williams, Christine Siddoway, Colin Shaw, Micah Jessup, Chris Andronicos, Matt Heizler, Joe Kopera, and Steve Rogers. Technical assistance and patient instruction by Kathy Manser in the U-Pb lab was essential to this project and made geochronology quite enjoyable. John Lansdown’s time and energy in assisting with the laser and Platform ICP-MS helped make detrital zircon geochronology possible. Adam Krawiec provided field assistance for one summer during this study. This study was primarily supported by National Science Foundation grant EAR 0003528 awarded to K. Karlstrom, M. Williams, J. Connelly, and C. Siddoway. Funding was also provided by the Department of Geological Sciences and the Geology Foundation at The University of Texas at Austin. The 2001 Keck Foundation summer research project led by Chris Siddoway in the Wet Mountains helped to develop a number of ideas presented in Chapter 3 and provided occasional food and fellowship. Finally, I want to thank my family and friends for support and encouragement throughout the course of my studies. Rebecca and Lucy have demonstrated exceptional patience, and none of this would have been possible without them. Onward!

iv

Proterozoic tectonic evolution of southern Laurentia: New constraints from field studies and geochronology in southern Colorado and northern New Mexico, U.S.A.

Publication No.______

James V. Jones III, Ph.D. The University of Texas at Austin, 2005

Supervisor: James N. Connelly

New field studies and geochronology from southern Colorado and northern New Mexico constrain the Proterozoic growth and modification of southern Laurentia. The Sangre de Cristo Mountains of southern Colorado preserve evidence for three episodes of Proterozoic magmatism, deformation, and metamorphism. Early deformation produced penetrative, NW-striking fabrics and occurred in an arc setting between 1750 – 1730 Ma.

Post-Yavapai granitoid magmatism occurred at 1695±2 Ma and 1682±3 Ma and was contemporaneous with regional quartzite deposition. Deformation at 1637±6 Ma produced localized NE-striking, subvertical fabrics with dextral shear sense. Granitic

magmatism at 1434±2 Ma and 1407±6 Ma was accompanied by NW – SE shortening between 1420 – 1412 Ma that produced subvertical, NE-striking fabrics.

v Thick sequences of quartz arenite were deposited across the region between the Yavapai and Mazatzal orogenies (ca. 1.70 Ga and 1.65 Ga). New geochronology reveals that deposition occurred on exhumed, Yavapai-aged basement (1706±5 Ma and 1698±4 Ma) with detritus dominated by Paleoproterozoic sources only slightly older than the quartzites themselves. Regional quartzite sedimentation was contemporaneous with nearly continuous magmatism in the region at deeper crustal levels. The first-cycle, syn- orogenic character of quartzites contrasts with their extreme compositional maturity, requiring perhaps anomalous environmental influences that enhanced chemical weathering during deposition.

New geochronology and structural studies from the Wet Mountains, Colorado, reveal contrasting structural styles during widespread Mesoproterozoic A-type granitic magmatism. At shallower crustal levels, strongly localized deformation at 1430+5/-3 Ma produced subvertical fabrics throughout the N-striking Five Points shear zone. At deeper crustal levels, penetrative deformation accompanying granitic magmatism at 1435±5 Ma and 1390±10 Ma produced moderately- to shallowly-dipping fabrics. Regionally consistent fabric orientations and kinematics are interpreted to represent an intracontinental response to convergent tectonism, and contrasting, yet coeval, styles of deformation require a structural discontinuity in the middle crust between ca. 1430 – 1360 Ma. Weak, flowing lower crust is consistent with models for intraplate orogenesis and the development of orogenic plateaus, and the southern Wet Mountains might represent an exhumed analog for mid-crustal, low-viscosity layers inferred beneath modern intracontinental orogenic systems such as Tibet and the Altiplano.

vi Table of Contents

List of Tables ...... ix

List of Figures...... x

Introduction...... 1

Chapter 1. Proterozoic tectonic evolution of the Sangre de Cristo Mountains, southern Colorado...... 4 Abstract...... 4 Introduction...... 5 Geologic Background ...... 7 Proterozoic Geology Of The Sangre De Cristo Mountains ...... 10 U-Pb Zircon And Titanite Geochronology ...... 29 Proterozoic Tectonic History Of The Sangre De Cristo Mountains...... 40 Regional (SW UW) Correlations...... 48 Regional (SW UW) Tectonic Implications...... 52 Summary And Conclusions ...... 56

Chapter 2. New age (U-Pb and Pb-Pb) constraints on the deposition and deformation of Proterozoic quartzites across the southern Rocky Mountains...... 63 Abstract...... 63 Introduction...... 64 Geologic Background ...... 67 Blue Ridge Quartzite, Colorado...... 71 Quartzite Detrital Zircon Geochronology...... 86 Discussion...... 94 Conclusions...... 103 Implications...... 104 Laurentian Correlative Sequences ...... 105 Summary...... 107

vii Chapter 3. Contrasting structural styles at ca. 1.4 Ga across the Wet Mountains, Colorado: Implications for models for intracontinental tectonism in the southern Rocky Mountains ...... 110 Abstract...... 110 Introduction...... 111 Geologic Setting...... 113 Proterozoic Lithologies And Structural Elements Of The Wet Mountains 116 U-Pb Zircon And Titanite Geochronology ...... 130 Discussion...... 152 Summary/Conclusions ...... 166

Appendix 1...... 169 Methods: Isotope Dilution Thermal Ionization Mass Spectrometer (ID-TIMS) analysis...... 169

Appendix 2...... 171 Methods: Laser Ablation inductively coupled mass spectrometer (LA-ICP-MS) analysis...... 171

Appendix 3...... 173 Quartzite LA-ICP-MS data...... 173

Bibliography ...... 186

Vita…...... 201

Back Plates...... 202

viii List of Tables

TABLE 1.1. SANGRE DE CRISTO MOUNTAINS U-PB ISOTOPIC DATA AND AGES. 59 - 61

TABLE 1.2. SUMMARY OF PROTEROZOIC TECTONIC EVENTS IN THE SANGRE DE CRISTO MOUNTAINS, COLORADO 62

TABLE 2.1. BLUE RIDGE AND QUARTZITE U-PB DATA AND AGES 81 - 82

TABLE 2.2. QUARTZITE DETRITAL ZIRCON 207PB/206PB AGE SUMMARY 88

TABLE 3.1. WET MOUNTAINS U-PB ISOTOPIC DATA AND AGES. 135 - 137

TABLE 3.2. SUMMARY OF U-PB AGES AND STRUCTURAL ELEMENTS IN THE WET MOUNTAINS, COLORADO 153 - 154

ix List of Figures

FIGURE 1.1. REGIONAL (SOUTHWESTERN U.S.) INDEX MAP. 9

FIGURE 1.2. GENERALIZED PRECAMBRIAN GEOLOGY OF THE SANGRE DE CRISTO MOUNTAINS, SOUTHERN COLORADO. 12

FIGURE 1.3. GEOLOGIC MAP OF THE MARSHALL GULCH AREA. 16

FIGURE 1.4. FIELD PHOTOGRAPHS FROM THE MARSHALL GULCH AREA. 17

FIGURE 1.5. GEOLOGIC MAP OF THE CRESTONE AREA. 20

FIGURE 1.6. FIELD PHOTOGRAPHS FROM THE MUSIC PASS AREA. 25 - 26

FIGURE 1.7. FIELD PHOTOGRAPHS OF PEGMATITE DIKES CUTTING THE SOUTHERN MARGIN OF THE MARSHALL GULCH PLUTON. 28

FIGURE 1.8. U-PB CONCORDIA AND ISOCHRON DIAGRAMS FOR SAMPLES FROM THE MARSHALL GULCH AREA. 31

FIGURE 1.9. U-PB CONCORDIA DIAGRAMS FOR SAMPLES FROM THE CRESTONE AREA. 36

FIGURE 1.10. U-PB CONCORDIA DIAGRAMS FOR SAMPLES FROM THE MUSIC PASS AREA. 39

FIGURE 1.11. SCHEMATIC BLOCK DIAGRAMS ILLUSTRATING THE TEMPORAL AND SPATIAL RELATIONSHIP BETWEEN PROTEROZOIC MID-CRUSTAL MAGMATISM AND DEFORMATION IN THE SANGRE DE CRISTO MOUNTAINS. 41

FIGURE 2.1. PALEOPROTEROZOIC QUARTZITES OF SOUTHWESTERN LAURENTIA. 65

FIGURE 2.2. PALEOPROTEROZOIC QUARTZITE EXPOSURES AND INFERRED REGIONAL EXTENT ACROSS COLORADO AND NEW MEXICO. 68

FIGURE 2.3. GENERALIZED GEOLOGIC MAP OF THE PROTEROZOIC QUARTZITE- SCHIST SEQUENCE EXPOSED ALONG BLUE RIDGE, COLORADO. 72

FIGURE 2.4. SCHEMATIC GEOLOGIC CROSS SECTIONS ACROSS GOOSEBERRY GULCH SYNCLINE, BLUE RIDGE, COLORADO. 73

x FIGURE 2.5. SCHEMATIC ILLUSTRATION OF CROSS-CUTTING RELATIONSHIPS AND SUMMARY OF NEW U-PB AND DETRITAL ZIRCON GEOCHRONOLOGY, BLUE RIDGE, COLORADO. 76

FIGURE 2.6. U-PB CONCORDIA DIAGRAMS FOR SAMPLES FROM BLUE RIDGE, COLORADO. 80

FIGURE 2.7. FIELD PHOTOGRAPH OF PEGMATITE DIKE CUTTING AND DEFLECTING QUARTZITE AND SCHIST FABRIC AT BLUE RIDGE, COLORADO. 85

FIGURE 2.8. DETRITAL ZIRCON AGE DISTRIBUTION PLOTS FOR EARLY PROTEROZOIC QUARTZITES FROM NORTHERN NEW MEXICO AND SOUTHERN COLORADO. 89 - 90

FIGURE 2.9. U-PB CONCORDIA DIAGRAMS FOR DETRITAL ZIRCON GRAINS FROM QUARTZITE SAMPLES. 91

FIGURE 3.1. REGIONAL (SOUTHWESTERN U.S.) INDEX MAP. 115

FIGURE 3.2. GENERALIZED PRECAMBRIAN GEOLOGY OF THE WET MOUNTAINS. 117

FIGURE 3.3. FIELD PHOTOGRAPHS FROM THE EASTERN ARKANSAS RIVER GORGE. 124

FIGURE 3.4. FIELD PHOTOGRAPHS OF THE CENTRAL AND SOUTHERN WET MOUNTAINS. 127

FIGURE 3.5. GENERALIZED GEOLOGIC MAP OF THE NORTH HARDSCRABBLE CREEK AREA, CENTRAL WET MOUNTAINS. 128

FIGURE 3.6. FOLIATION MAP FOR PRECAMBRIAN EXPOSURES IN THE BEAR CREEK/WILLIAMS CREEK AREA OF THE SOUTHERN WET MOUNTAINS. 129

FIGURE 3.7. U-PB CONCORDIA DIAGRAMS FOR SAMPLES FROM THE ARKANSAS RIVER GORGE, NORTHERN WET MOUNTAINS. 134

FIGURE 3.8. FIELD PHOTOGRAPHS FROM THE RATTLESNAKE GULCH AREA, CENTRAL WET MOUNTAINS. 140

FIGURE 3.9. U-PB CONCORDIA DIAGRAMS FOR SAMPLES FROM THE RATTLESNAKE GULCH AREA, CENTRAL WET MOUNTAINS. 141

FIGURE 3.10. U-PB CONCORDIA DIAGRAMS FOR SAMPLES FROM THE BEAR CREEK/WILLIAMS CREEK AREA, SOUTHERN WET MOUNTAINS. 146 xi FIGURE 3.11. FIELD PHOTOGRAPHS FROM THE BEAR CREEK/WILLIAMS CREEK AREA, SOUTHERN WET MOUNTAINS. 150 - 151

FIGURE 3.12. SCHEMATIC BLOCK DIAGRAM OF THE MIDDLE CRUST BENEATH SOUTHERN COLORADO AT CA. 1.4 GA. 162

BACK PLATES (NOTE: OVERSIZED COLOR PLATES ARE FORMATTED TO PRINT ON TABLOID-SIZED (11"X17") PAPER) 202

PLATE 1. GEOLOGIC MAP OF THE MUSIC PASS AREA, SANGRE DE CRISTO MOUNTAINS, COLORADO. 2023

PLATE 2. GENERALIZED GEOLOGIC MAP OF THE EASTERN ARKANSAS RIVER GORGE (ARG), WET MOUNTAINS, COLORADO. 2024

xii Introduction

Laurentia, the Precambrian cratonic core underlying much of North America, was first formed through the aggregation of Archean microcontinents across Proterozoic mobile belts between 2.0 – 1.8 Ga (Hoffman, 1988, and references therein). Following this early stage of assembly, an active margin developed across southern Laurentia, and a belt of juvenile crust more than 1000 km wide was accreted between 1.80 – 1.65 Ga. Much of the Precambrian continental lithosphere of the southwestern United States formed during this time. After a period of relative tectonic quiescence, southern Laurentia was intruded by a voluminous suite of A-type granitic magmas at ca. 1.4 Ga, possibly related to continued southward growth of the continent between 1.5 – 1.3 Ga. The southward growth of Laurentia continued between 1.3 – 1.0 Ga and culminated with the ca. 1.1 Ga Grenville Orogeny and assembly of the supercontinent Rodinia. Although models for the southward growth of Laurentia and evolution of the Proterozoic continental lithosphere are generally widely accepted (e.g., Condie, 1982; Bowring and Karlstrom, 1990), numerous complexities remain to be resolved. These problems include the following: 1) the age and tectonic setting of early, NW-striking fabrics throughout the southwestern U.S. that are essentially orthogonal to the regional, NE-striking tectonic grain (e.g., Jessup et al., 2005); 2) the age and tectonic/depositional setting of widespread, thick sequences of supermature quartzite along the southern Laurentian margin (e.g.,

1 Soegaard and Eriksson, 1989); and 3) the tectonic setting and significance of widespread, voluminous A-type granitic magmatism at ca. 1.4 Ga (e.g., Anderson, 1983; Nyman et al., 1984). This study presents new field observations and geochronology (U-Pb and Pb-Pb) from Precambrian exposures throughout southern Colorado and northern New Mexico that provide new insight into these tectonic problems and further constrain existing regional models for the Proterozoic tectonic evolution of southern Laurentia. The following three chapters were written as individual manuscripts that will be submitted to peer-reviewed journals. Chapter 1 presents new field mapping and results from the Sangre de Cristo Mountains of southern Colorado that constrain the tectonic evolution of the range from early arc magmatism, deformation, and metamorphism during the Paleoproterozoic to newly-discovered episodes of granitic magmatism, deformation, and metamorphism during the Mesoproterozoic (ca. 1.4 Ga). Chapter 2 presents new geochronology from Proterozoic quartzite exposures across southern Colorado and northern New Mexico that constrains the age of regional quartzite deposition. Detrital zircon ages also provide new information regarding the sedimentary provenance and detrital character of the metasedimentary sequences. Chapter 3 focuses on the Wet Mountains of southern Colorado, and new structural observations and geochronology from N – S across the range reveal new evidence for a ca. 1.4 Ga structural discontinuity in the middle crust separating localized, subvertical fabrics in the north from widespread, penetrative deformation throughout the central and southern part of the range. These contrasting structural styles

2 accompanied widespread granitic magmatism and have important implications for regional tectonic models for ca. 1.4 Ga magmatism and for processes involved during intracontinental orogenesis.

3 Chapter 1. Proterozoic tectonic evolution of the Sangre de Cristo Mountains, southern Colorado

ABSTRACT

Field studies and U-Pb geochronology from the Sangre de Cristo Mountains, southern Colorado, provide new constraints on the Proterozoic tectonic evolution of southern Laurentia. Early deformation and metamorphism was broadly coeval with arc-related magmatism between ca. 1750 – 1720 Ma and produced gneissic fabrics that are subvertical and strike NW. These NW-striking fabrics are interpreted to have formed during early, NE – SW convergent tectonism, perhaps in outboard arcs, prior to northward accretion to the southern margin of Laurentia. Whereas early deformation was penetrative in character, subsequent Proterozoic deformation was strongly localized throughout the range. Emplacement of granitic and tonalitic plutons occurred between ca. 1695 – 1682 Ma and was coeval with widespread deposition of thick sequences of supermature quartzite. Magmatism and quartzite sedimentation are interpreted to result from contemporaneous crustal responses to collapse or relaxation of the Yavapai accretionary orogen between ca. 1700 – 1660 Ma. Deformation related to the Mazatzal Orogeny occurred at 1637±6 Ma locally, involving dextral shear in NE- trending shear zones and producing subhorizontal, NW – SE shortening features such as subvertical, NE-striking foliation and large scale, upright folding of older layering. Quartzite is inferred to have been deformed during this time and buried

4 to granitic emplacement depths (10 – 15 km). Two newly-dated Mesoproterozoic granitic intrusions were emplaced at 1434±2 Ma and 1407±6 Ma, respectively, accompanied by development of a NE-striking, subvertical tectonic foliation and localized shear zones between ca. 1420 – 1412 Ma. The map-scale geometry of these intrusions and coeval deformation suggest that ca. 1.4 Ga granites were emplaced into a broadly convergent strain field, and the inhomogeneous character of deformation is consistent with models for NW – SE shortening deformation throughout the region.

INTRODUCTION

Although regional tectonic models for the Proterozoic growth and evolution of southern Laurentia are generally widely accepted, some significant local and regional complexities remain to be resolved. First, numerous models have been proposed to explain early, NW-striking fabrics across the southwestern U.S. that predate and are essentially perpendicular to regional NE-striking fabrics related to accretion of the Yavapai and Mazatzal crustal provinces (Karlstrom, 1989; Duebendorfer et al., 2001; Jessup et al., 2005). Second, cross-cutting relationships and structural arguments suggest that locally-thick (1 – 2 km), supermature quartzite sequences were deposited across a broad region during a ca. 50 m.y. break in Paleoproterozoic accretionary orogenesis (Williams et al., 2003), but absolute age constraints on their deposition and deformation are generally lacking. Third, the tectonic setting of widespread A-type granitic magmatism at ca. 1.4 Ga has been a longstanding subject of debate. Whereas granites were

5 historically considered to be “anorogenic” based on their distinct geochemistry and a perceived lack of contemporaneous deformation (Anderson, 1983), more recent work has suggested that ca. 1.4 Ga granites were emplaced during a regional intracontinental orogenic event (Graubard and Mattison, 1990; Nyman et al., 1994). New observations and data bearing on the resolution of these various controversies not only have important implications for constraining local and regional tectonic histories, but they also directly bear on the collective interpretation of fundamental tectonic processes governing the growth, modification, and stabilization of Laurentian continental lithosphere throughout the Proterozoic. Exposures of basement gneiss, quartzite, and granitic intrusive rocks throughout the Sangre de Cristo Mountains provide a relatively complete record of the Proterozoic tectonic evolution of southern Colorado. This study combines new, detailed field mapping of previously undated plutons with precise U-Pb geochronology in the central part of the range to constrain the timing of magmatism, metamorphism, and the development and reactivation of Proterozoic tectonic fabrics. Combined with published geochronology from other parts of the mountain range, new results illustrate the episodic nature of igneous and tectonic activity during the Proterozoic and provide new constraints on regionally recognized tectonic events. These results constrain an important change from early NW-striking, penetrative fabrics to localized, NE-striking tectonic fabrics that characterize Proterozoic exposures throughout the southern Rocky Mountains. Furthermore, new structural observations suggest that pre-existing

6 structural weaknesses and/or local anisotropies fundamentally controlled the geometry and spatial distribution deformation and magmatism throughout the Proterozoic and Phanerozoic. Finally, geochronologic data reveal two previously unrecognized pulses of Mesoproterozoic (ca. 1.4 Ga) granitic magmatism that, in part, provide a new minimum age on the deposition of quartzite in this region. The geometry of plutons and the orientation of fabrics developed at ca. 1.4 Ga suggest that the regional shortening direction was consistently oriented over ca. 30 m.y. However, exposures record spatially and temporally heterogeneous deformation that is consistent with the general structural style of ca. 1.4 Ga deformation at shallower mid-crustal levels throughout the southern Rocky Mountains.

GEOLOGIC BACKGROUND

Precambrian crustal exposures across the southwestern U.S. comprise a diverse compositional assemblage of metavolcanic rocks, metasedimentary rocks, and mafic and granitoid plutons that were accreted to the southern margin of the Archean Wyoming Province between 1.8 – 1.6 Ga (Condie, 1982; Karlstrom and

Bowring, 1988; Reed et al., 1993) as part of a protracted period of Laurentian crustal growth. The Yavapai Province represents the earliest phase of Paleoproterozoic crustal growth and accretion during which a complex collage of juvenile arc terranes were added between 1.78 – 1.70 Ga along a belt stretching from Colorado to Arizona. This phase of arc growth and collision culminated with the ca. 1.71 – 1.70 Ga Yavapai Orogeny (Karlstrom and Bowring, 1988).

7 Following a prolonged episode (ca. 40 m.y.) of voluminous post-orogenic granitoid magmatism and regional quartzite deposition (Anderson and Cullers, 1999; Williams et al., 2003), the Mazatzal Province was accreted to the south along a belt extending across New Mexico and Arizona during the 1.66 – 1.62 Ga Mazatzal Orogeny (Silver, 1965; Karlstrom and Bowring, 1988). After a ca. 200 my tectonic lull, renewed southward growth of Laurentia is inferred to have occurred during the Mesoproterozoic, based on a large crustal province with Nd model ages of 1.5 – 1.3 Ga extending from northern Mexico to Labrador (Bennett and DePaolo, 1987; Patchett and Ruiz, 1989; Karlstrom et al., 2001). A widespread episode of granitic magmatism, local mafic diking, and regional high-temperature, low-pressure metamorphism occurred throughout the southwestern U.S. between 1.47 – 1.36 Ga (Reed et al., 1993; Williams, 1993). Granites of this age account for nearly 20% of all Precambrian exposures across the Rocky Mountain region and were widely emplaced into the newly accreted Paleoproterozoic crust but are not found among Archean exposures to the north. Ca. 1.4 Ga granites are characterized by distinct, A-type geochemical characteristics (Loiselle and Wones, 1979; Anderson, 1983) that are commonly associated with extensional tectonic environments (e.g., continental rifting; Emslie, 1978). However, regional evidence for deformation within the thermal aureoles of plutons and contemporaneous reactivation of NE-striking crustal shear zones suggests that magmatism throughout the southwestern U.S. was accompanied by regional NW – SE shortening at ca. 1.4 Ga (Graubard and Mattinson, 1990; Shaw et al., 2001).

8

Figure 1.1. Regional (southwestern U.S.) index map. Precambrian exposures (grey) and ca. 1.4 Ga granites (red) are emphasized. Proterozoic crustal provinces, inferred boundaries and/or transition zones, and approximate age ranges are shown (Condie, 1986; Bennett and DePaolo, 1987; Karlstrom and Bowring, 1988; Wooden et al., 1988; Wooden and DeWitt, 1991). Regional qualitative strain ellipse inferred from structural data from the southern Rocky Mountains and Arizona (Graubard and Mattison, 1990; Kirby et al., 1995; Nyman and Karlstrom, 1997; Shaw et al., 2001; Selverstone et al., 2000).

9 During the southward Paleoproterozoic growth of Laurentia, thick (1 – 2 km) sequences of quartz arenite were deposited along the southern margin of the continent. Some of the best examples of these deposits across the southwestern U.S. include large, extensively exposed sequences (e.g., Uncompahgre Formation, Ortega Quartzite) as well as numerous smaller, localized exposures throughout the region. Quartzites in the metasedimentary sequences are characterized by a high degree of compositional maturity and preserved primary sedimentary structures, and they overlie polydeformed basement assemblages comprising compositionally diverse, but commonly mafic, lithologies (e.g., Dubois and Cochetopa successions; Bickford and Boardman, 1984). After deposition, both quartzite sequences and underlying basement assemblages were deformed, and exposures across southern Colorado are commonly preserved as synclinal “keels” interpreted to represent the roots of much larger folds that have been eroded away (Reuss, 1974). Whereas published age constraints (U-Pb zircon) from Arizona and New Mexico require that quartzite deposition occurred during the ca. 1.7 Ga Yavapai Orogeny (1703 Ma, Cox et al., 2002b; Bauer and Williams, 1989), the minimum age of quartzite deposition in southern Colorado is poorly constrained because cross-cutting igneous bodies generally yield ages that are too young (i.e., ca. 1.4 Ga) to narrowly define the age of deformation.

PROTEROZOIC GEOLOGY OF THE SANGRE DE CRISTO MOUNTAINS

The Sangre de Cristo Mountains, southern Colorado, lie in the southern part of the Yavapai Province within the broad zone affected by Mazatzal-aged 10 orogenesis (Fig. 1.1). The range is also located within a region including southern Colorado and New Mexico that preserve evidence of widespread deposition of quartzite during and after the ca. 1.7 Ga Yavapai Orogeny and extensive granitic magmatism at ca. 1.4 Ga. Although Proterozoic exposures throughout the Sangre de Cristo Mountains are locally segmented by Phanerozoic reverse and normal faults (Fig. 1.2), reconnaissance mapping (Jones, unpublished mapping) suggests that there is structural continuity between individual thrust sheets along the length of the range. Proterozoic exposures comprise a compositionally diverse assemblage of felsic to mafic metavolcanic rocks, interlayered metasedimentary rocks, and local cross-bedded quartzite that is intruded by voluminous mafic to granitic rocks of both Paleoproterozoic and Mesoproterozoic age. New field mapping reveals that these rocks preserve evidence for at least three episodes of deformation and metamorphism (D1 – D3 and M1 – M3, respectively) and record an important change in the orientation of regional deformational fabrics from NW to NE strikes between D1 and D2. Fabrics developed during D1 and M1 are preserved across the full length of the range (>100 km) and influenced the geometry and localization of subsequent tectonism ranging from Late Paleo- to Mesoproterozoic magmatism and deformation to multiple phases of Phanerozoic brittle deformation. In contrast, younger Proterozoic deformation (D2 and D3) produced fabrics that are similar in orientation and are parallel with a well- documented, regional NE – SW tectonic grain, but these fabrics are strongly localized into discrete domains of concentrated deformation throughout the range.

11

Figure 1.2. Generalized Precambrian geology of the Sangre de Cristo Mountains, southern Colorado, and summary of new and existing U-Pb geochronology. New zircon (z) and titanite (t) ages from this study are described in text. Previously published ages are all U-Pb zircon, and the respective data sources are indicated by asterisks. Average foliation (poles to S1) orientation is represented for three parts of the range on lower-hemisphere, equal-area stereonet diagrams. Structural data in the central part of the range were compiled from this study and Johnson et al. (1987). Structural data in the northern and southern Sangre de Cristo Mountains were compiled from Johnson et al. (1987). Structural analysis was done by the author using GEOrient 9.1 (Holcombe, 2003).

12 The following sections describe the various Proterozoic rocks exposed in the Sangre de Cristo Mountains, organized from oldest to youngest. The dominant lithologies are discussed in conjunction with the various structural elements they contain because these were the criteria used to select those rocks best suited for new U-Pb geochronology to constrain the Proterozoic tectonic evolution of the range.

Mixed gneiss and amphibolite

The oldest rocks exposed in the Sangre de Cristo Mountains are a sequence of fine- to medium-grained gneisses and amphibolite that make up the host rock to all intrusive phases described below. These rocks, along with Paleoproterozoic-aged intrusive rocks described below, are also the local basement underlying a thin band of cross-bedded quartzite in the central part of the range (see below). Whereas these rocks comprise a highly-variable assemblage described in detail by Johnson et al. (1987), there is a general compositional gradation from more felsic gneisses and schistose rocks in the northern part of the range (southernmost contact at northern end of Marshall Gulch pluton; Fig. 1.3) to interlayered, salt-and-pepper felsic and mafic gneisses in the central part of the range to more mafic gneisses and amphibolite in the southern part of the range. These various compositions are commonly interlayered with one another and have contacts that range from sharp to gradational across tens of centimeters.

13 Structurally, gneissic rocks across the range are characterized by a subvertical foliation that strikes NW – SE and is especially well developed in the northern and southern parts of the range (Fig. 1.2). Foliations in the central part of the range exhibit a departure from this dominant fabric and, instead, strike ENE and dip steeply north. Reconnaissance mapping by the author suggests that these N-dipping foliations are more common on the western side of the range where basement gneisses contain E-striking fabrics that are developed within localized high-strain domains. However, many of these rocks are difficult to access, and their tectonic significance has yet to be determined. Whereas compositional domains that likely reflect primary layering define the dominant foliation across the range, individual compositional layers are locally folded by centimeter-scale isoclinal folds (Fig. 1.4A). Foliation is axial planar to these folds, and, thus, the dominant gneissic fabric is interpreted to represent a minimum of one early deformational event (S1).

Blanca Peak intrusive suite

A suite of intermediate and mafic intrusive rocks is exposed at the southern end of the range south of Medano Pass in the Blanca Peak area (Fig. 1.2). These medium- to coarse-grained plutonic rocks are described in detail by Sabin (1994, after Johnson et al., 1987) and include (in decreasing volumetric order) tonalite, diorite, granodiorite, quartz diorite, and gabbro. These rocks intrude amphibolite-facies mafic and felsic gneisses and amphibolite and cut across but also locally deflect the NW-striking gneissic foliation (S1). The

14 plutonic rocks locally contain a solid-state foliation, but reconnaissance mapping indicates that they are not strongly deformed. Sabin (1994) dated a suite of Blanca Peak intrusive rocks and identified two age populations, an older group with ages of ca. 1750 Ma and a younger group with ages of ca. 1730 Ma (Fig. 1.2). These rocks were not the focus of this study, but they provide useful age constraints on the deposition and/or crystallization of basement gneisses and development of the early NW-striking fabric across the range.

Marshall Gulch syenite and monzogranite (Xsm).

The Marshall Gulch pluton is exposed on the western flank of the Sangre de Cristo Mountains, ca. 10 km north of the town of Crestone (Fig. 1.2) between the Rito Alto and Wild Cherry Creek U.S. Forest Service trailheads. White to gray syenite and monzogranite are the main compositional phases of this granitic body, and modal mineral percentages are approximately 50 – 80% microcline, as much as 35% plagioclase, 1 – 5% quartz, as much as 10% biotite, minor muscovite, and accessory titanite, zircon, and apatite (Johnson et al., 1987). Exposures are commonly coarse-grained to K-feldspar megacrystic with individual microcline phenocrysts up to 4 cm wide and 8 cm long, but locally there are occurrences of much finer-grained, leucocratic granite. Both of these phases contain the same solid-state foliation, and the contact between them is gradational and commonly interfingering, such that they are interpreted to be part of the same general intrusive event. Although it was undated, previous workers

15

Figure 1.3. Geologic map of the Marshall Gulch area. Contacts were modified from Lindsey et al. (1985) and Lindsey et al. (1986). Lower-hemisphere, equal-area stereonet diagrams represent poles to foliation in the Marshall Gulch monzogranite (A) and host-rock gneiss (B) and lineation data from granitoid rocks (C). All structural data and analysis are from this study. Interpreted fabric elements were determined primarily based on orientation and are discussed in the text along with evidence for relative timing relationships. Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003). Arrows indicate approximate kinematic sense recorded by particular fabric elements. 16

Figure 1.4. Field photographs from the Marshall Gulch area. A) View down on small isoclinal fold hook in host-rock gneiss ca. 10 m from the contact with Marshall Gulch monzogranite. B) View down on deformed Marshall Gulch monzogranite. Granitoid matrix has undergone grain size reduction, and K- feldspar porphyroclasts have asymmetric, recrystallized tails that indicate dextral offset. C) Tight, upright fold affecting monzogranite and host-rock amphibolite. Note boudinage of hinge zone. View to NE. D) View down on asymmetric mineral fabric (S-C) in coarse-grained monzogranite indicating dextral shear sense.

17 correlated the Marshall Gulch pluton with a regional suite of Paleoproterozoic granitic intrusions based on its coarse-grained texture and the presence of deformational fabrics (Routt plutonic suite of Tweto, 1987; Tweto, 1979; Johnson et al., 1987). New U-Pb geochronology results described below support this correlation. Host rocks to the Marshall Gulch pluton generally consist of upper greenschist- to amphibolite-facies gneiss, schist, and amphibolite, but there is a distinct change from more felsic, mica-rich to schistose gneisses in the north to more mafic, salt-and-pepper gneisses and amphibolite to the south. In the southern part of the pluton, a large (ca. 0.5 km diameter) composite xenolith contains medium- to coarse-grained tonalite that interfingers with the gneisses and is cut sharply by the granitic rocks. These host rocks contain a well-developed foliation that strikes NNW and dips steeply WSW (Fig. 1.3B). This foliation is interpreted to be a polyphase/composite fabric (S1) based on local isoclinal folding of the dominant foliation (Fig. 1.4A). The pluton itself was emplaced parallel to the host-rock fabric as an elongate, tabular body (2 km wide by 7 km long as exposed) and contains a well-developed magmatic foliation defined by aligned tabular feldspars and biotite that is parallel with foliation in the surrounding gneiss (average strike/dip=330/70°SW, Fig. 1.3A). This magmatic fabric (Smagmatic) is present throughout the pluton, but it is best developed within 0.5 km of the intrusion margin. Across the southernmost 1 km of its exposure, the granite contains a solid-state foliation (S2) that is oriented essentially perpendicular to the magmatic and overall host-rock fabric (Fig. 1.3A). This

18 well-developed foliation is also contained in host-rock gneiss and tonalite in the southern part of the composite xenolith (Fig. 1.3B). In granitic rocks, it is primarily defined by biotite and is enhanced locally by grain-size reduction of the originally coarse-grained granitic matrix (Fig. 1.4B and D). The orientation of S2 is parallel with axial planes of tight, upright folds in thin (20 – 30 cm) dikes of granitic rock (Fig. 1.4C). Where microcline megacrysts are locally preserved, they have asymmetric recrystallized tails indicating dextral offset along a shallowly plunging biotite mineral lineation (Fig. 1.4B). In outcrops where grain size reduction has not occurred, coarse-grained granite locally contains a composite (S-C) foliation recording dextral shear sense (Fig. 1.4D). The presence of both shortening- and shear-related deformation features associated with the development of S2 suggests that the Marshall Gulch pluton experienced at least one episode of kilometer-scale NW – SE directed shortening accompanied by perhaps tens of meters of dextral shear after crystallization.

Crestone quartz monzonite (Xqm)

The Crestone stock is exposed on the western flank of the Sangre de Cristo Mountains and forms the prominent, jagged ridge just ENE of the town of Crestone (Fig. 1.2). This intrusion comprises light gray to light brown to tan, fine- to medium-grained quartz monzonite with modal percentages estimated at 35% plagioclase, 30% microcline, 30% quartz, 2 – 5% biotite, minor muscovite, and accessory magnetite, titanite, zircon, and apatite (Johnson et al., 1987). It is exposed as an elongate (2 km wide by 5 km long as exposed) body but is bounded

19

Figure 1.5. Geologic map of the Crestone area. Contacts modified from Lindsey et al. (1986). Lower-hemisphere, equal-area stereonet diagrams represent poles to foliation in the Crestone quartz monzonite (A) and in host-rock gneiss (B) and lineation data from granitoid rocks (C). All structural data and analysis are from this study. Interpreted fabric elements were determined primarily based on orientation and are discussed in the text along with evidence for relative timing relationships. Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003). 20 by Phanerozoic thrust faults on two sides (Fig. 1.5). It intrudes greenschist-facies amphibolite and fine-grained salt-and-pepper gneiss that are only exposed along its southern end and as small (<1 m) xenoliths in the heart of the intrusion. The granite cuts a well-developed, pre-existing gneissic foliation (S1) in host-rock gneisses, and the contact is very sharp. Away from the southern contact, the granite is essentially undeformed with the exception of a locally developed mica foliation that does not display a consistent orientation. The apparent lack of deformation across much of the Crestone stock and its fine-grained nature led previous workers to correlate this intrusion with a regional suite of granitic intrusions that are Mesoproterozoic in age (Berthoud plutonic suite of Tweto, 1987; Tweto, 1979; Johnson et al., 1987), but it was previously undated. New U- Pb geochronology reported below indicates that it was emplaced during the Paleoproterozoic. The southernmost exposures of the granite contain a well-developed foliation defined by biotite and elongate quartz grains that strikes NE – SW and dips steeply SE (Fig. 1.5A). This foliation is parallel with the dominant foliation in host-rock gneiss and amphibolite within 0.25 km of the margin of the stock

(Fig. 1.5B), but the host-rock foliation changes to a more N – S orientation away from the granite to the south. The solid-state foliation is similar in orientation to both S2 of the Marshall Gulch pluton and S3 of the Music Pass pluton (described below), but timing relationships discussed below suggest that the foliation formed or was at least enhanced during Mesoproterozoic deformation (D3), and is, therefore, an S3 fabric.

21 Quartzite (Xq)

In the central part of the range, a 100 m thick layer of cross-bedded quartzite is exposed along the eastern margin of the Music Pass pluton (Plate 1). Quartzite is variable in appearance with colors ranging from white to gray to red, is relatively pure (90% quartz), and contains relict layering defined by seams rich in opaque minerals. The nature of the contact between the quartzite and underlying basement is difficult to determine as it is only locally exposed 0.5 km southeast of Snowslide Mountain (Plate 1) and is deeply weathered. However, based on subtle contrasts in the orientation of compositional layering within 10 meters of the contact, the quartzite is interpreted to be in either fault or unconformable depositional contact with underlying basement. Whereas gneissic rocks 15 meters from the contact also contain abundant, centimeter-sized isoclinal fold hooks, the quartzite does not appear in outcrop to be strongly deformed. Although the underlying basement assemblage is interpreted to include both gneiss and Paleoproterozoic intrusive rocks described above, the quartzite is only directly intruded by the Mesoproterozoic Music Pass pluton (see below). Based on new constraints from correlative quartzite sequences 70 km to the north (Blue Ridge, Colorado; see Chapter 2 of this study) and regional structural and geochronologic arguments, quartzite deposition might have been broadly coeval with Paleoproterozoic granitic magmatism in the Sangre de Cristo Mountains. This interpretation is discussed in greater detail below.

22 Music Pass quartz monzonite (Yqm)

The Music Pass pluton is exposed on the eastern flank of the Sangre de Cristo Mountains 20 km south of the town of Westcliffe (Fig. 1.2). Whereas the bulk of this pluton and its western margin are best exposed in the peaks and ridges west of Sand Creek (Plate 1), access to the mouth of the creek from the west side of the range is extremely difficult. Instead, exposures in this part of the range are best accessed from the east over Music Pass. The Music Pass pluton is characterized by grey to pink, coarse-grained to K-feldspar megacrystic quartz monzonite (Fig. 1.6D). Large pink to white microcline megacrysts up to 6 cm in length make up 25 – 45% of the rock, and estimated modal percentages of the groundmass are 60 % plagioclase, 20 % quartz, 10 – 20 % biotite and amphibole, up to 1 % titanite and magnetite, and accessory zircon and apatite (Johnson et al., 1987). The quartz monzonite was emplaced into amphibolite-facies medium- to coarse-grained gneiss, amphibolite, and quartzite along a well-developed, pre- existing foliation striking WNW – ESE with a subvertical dip (S1, Plate 1A). Gneisses surrounding the Music Pass pluton locally contain abundant, small (10 – 15 cm) isoclinal folds that affect both compositional layering and the gneissic foliation, indicating a polyphase deformational history prior to emplacement of the granite (Fig. 1.6A). Locally, the granite cuts across the pre-existing gneissic foliation of the host rocks, but over a scale of 3 – 5 m the granitic rocks and host rocks are intricately intermingled (Fig. 1.6B). On the scale of the pluton, a vertical, sharp contact is observed in the SE face of Tijeras Peak (Fig. 1.6C). The pluton is xenolith-poor except along its eastern margin around Snowslide

23 Mountain (Plate 1) where it contains large (up to 10 – 15 m2) blocks of quartzite. Based on its coarse-grained texture and widespread evidence for solid-state deformation within the pluton (see below), the Music Pass quartz monzonite was correlated by previous workers with a regional suite of Paleoproterozoic granitic intrusions (Routt plutonic suite of Tweto, 1987; Tweto, 1979; Johnson et al., 1987). Thus, it was sampled for U-Pb geochronology to determine the minimum depositional age of the host-rock quartzite. New results described below indicate that the pluton was emplaced during the Mesoproterozoic (ca. 1.4 Ga). The pluton contains a weakly- to moderately-developed foliation defined by large, tabular microcline megacrysts and biotite that is strongest within 0.5 – 1.0 km of the pluton margin. This fabric is interpreted to be magmatic in origin

(Smagmatic), and it strikes WNW – ESE and dips steeply, parallel to the contact and host-rock fabric. Detailed mapping within the pluton reveals that discrete zones of solid-state deformation up to a few meters thick also locally occur. These deformation zones are characterized by grain-size reduction of the coarse-grained granite matrix and development of a strong biotite foliation (S3) that strikes NE – SW and dips steeply NW. The general lack of asymmetry across many of these zones of deformation suggests that NW – SE shortening dominated during deformation, but local asymmetric fabrics, offset pegmatite dikes, and local composite (S-C) fabrics in granitic rocks suggest that some component of both sinistral and dextral shear locally accompanied shortening (Fig. 1.6F). One particular zone of deformation along the ridge SE of Music Pass displays mylonitic fabrics in which the granitic matrix has been reduced to sub-millimeter

24

Figure 1.6. Field photographs from the Music Pass area. A) View down on small isoclinal folds parallel to S1 in host-rock gneisses adjacent to the Music Pass pluton. B) Outcrop view of intricately intermingled character of contact between Music Pass quartz monzonite and host-rock gneiss. C) View to NW of contact between quartz monzonite and gneiss in the near-vertical face of Tijeras Peak. At this scale, the contact is relatively sharply defined and is subvertical, parallel to S1. 25

Figure 1.6 continued. D) K-feldspar megacrystic Music Pass quartz monzonite. Note magmatic fabric defined by aligned K-feldspar megacrysts. E) View down on mylonitic Music Pass quartz monzonite. F) View down on deformed Music Pass quartz monzonite containing asymmetric mineral fabrics defined by biotite and recrystallized K-feldspar megacrysts. Shear sense is sinistral.

26 grain sizes, and microcline megacrysts are elongated into ribbons with aspect ratios of up to 20:1 (Fig. 1.6E). Although the orientation and deformational styles associated with S3 correlate with the solid-state foliation contained within the southern margin of the Marshall Gulch pluton (S2), new U-Pb geochronology reported below indicates that the two fabrics formed at different times.

Pegmatite dikes (Yp)

Whereas each of the granitic plutons described above has at least one generation of pegmatite dikes and sills associated with its emplacement, the Marshall Gulch pluton is cut by a group of pegmatite dikes that do not appear to be directly related. These dikes intrude both basement gneisses and granitic rocks, occurring in a swarm of approximately 10 large intrusions with an average strike of 335° and near-vertical contacts (Figs. 1.3, 1.7). The most distinctive attribute of these dikes is their average thickness of 10 – 15 meters. These pegmatites are very leucocratic (white to light gray) and are composed of K- feldspar and quartz with minor muscovite, plagioclase, and black tourmaline. Contacts between the pegmatite and host rocks are commonly very sharp and planar to angular, suggesting that the host rocks were cool relative to the intruding material. Locally, angular blocks of host gneiss up to a meter in size are incorporated into the margins of the dikes, and smaller swarms of veins and dikes extend up to a few meters beyond the contacts of the larger dikes. Although previously undated, these dikes were interpreted to be Paleoproterozoic in age because they were emplaced into a Paleoproterozoic pluton (Lindsey et al., 1985).

27

Figure 1.7. Field photographs of pegmatite dikes cutting the southern margin of the Marshall Gulch pluton. View to NE.

28 However, new geochronology reported below indicates that they were emplaced during the Mesoproterozoic.

Mafic dike (Ym?)

The Music Pass pluton is cut on its northern end by a thick (30 m), subvertical mafic dike. The NNW-striking dike is primarily composed of coarse- grained gabbro and has no detectable internal fabric. It was sampled at one site for U-Pb geochronology but did not yield any mineral phases suitable for dating. However, it is correlated with a swarm of similar mafic dikes with NNW trends exposed in the Sawatch Range, Park Range, and northern Front Range of Colorado (Tweto, 1987, and references therein). These mafic intrusions are interpreted to be broadly synchronous with ca. 1.4 Ga granitic magmatism based on hornblende K-Ar ages and spatial and cross-cutting relationships with dated plutons (Tweto, 1987, and references therein).

U-PB ZIRCON AND TITANITE GEOCHRONOLOGY

The three granitic plutons described above, selected host rocks, and cross- cutting pegmatite dikes were sampled for U-Pb geochronology to constrain the age of magmatism, metamorphism, and deformation in the central Sangre de Cristo Mountains. Isotopic data are presented in Table 1.1, and associated concordia and isochron diagrams are presented in Figures 1.8 – 1.10. Zircon and

29 titanite fractions were hand picked, examined using a petrographic microscope, characterized by cathodoluminesence, extensively abraded, and then subjected to a final optical re-evaluation before analysis. For complete analytical methods refer to Appendix 1. New results described in the following section are grouped according to the map area from which samples were collected.

Marshall Gulch area

MARSHALL GULCH PLUTON (J01-MG1). A sample of foliated, K-feldspar megacrystic monzogranite from the southern part of the Marshall Gulch pluton (Fig. 1.3 for location) yielded a single population of large (0.5 mm on side), clear, blocky zircon interpreted to be fragments of larger zircon grains that were mechanically broken during mineral separation processes. Although no larger, whole zircon grains were recovered from coarser-grained sieve fractions, individual zircons up to 5 mm in length are visible in thin section. Two fractions

(Z1 and Z4) overlap concordia (Fig. 1.8A) and have an average 207Pb/206Pb age of 1695±2 Ma (Table 1.1). This age is interpreted to represent the time of crystallization of this pluton. Abundant pale yellow to clear fragments of titanite were also recovered from this sample. In thin section, titanite is observed to occur in clusters parallel to the solid-state biotite foliation. Three fractions (T1, T2, and T4) define a 238U/204Pb-206Pb/204Pb isochron with an age of 1637±6 Ma (MSWD=0.86, Fig. 1.8B). This age is significantly younger than the crystallization age of the monzogranite, and, thus, it could represent thermal resetting, growth of titanite

30

Figure 1.8. U-Pb concordia and isochron diagrams for samples from the Marshall Gulch area. Ages are defined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

31 during a post-crystallization metamorphic event, or metamorphic recrystallization and resetting of existing titanite during a metamorphism. Quartz and feldspar microtextures in the surrounding rock suggest that metamorphic conditions never exceeded lower- to middle-amphibolite facies during deformation of the Marshall Gulch monzogranite. These textures include brittle behavior of feldspar megacrysts together with dominantly rotational recrystallization of interstitial quartz grains. Temperatures associated with these metamorphic facies are generally thought to be lower than the U-Pb closure temperature for titanite

(~700°C; Pidgeon et al., 1996; Verts et al., 1996). Because titanite is believed to react readily during metamorphism (Frost et al., 2000), this age is interpreted to reflect metamorphic recrystallization of titanite, likely during deformation resulting in development of the solid-state biotite foliation (S2) across the southern margin of the pluton.

FOLIATED TONALITE (J01-MG2). A sample of medium-grained, foliated tonalite was collected from a large (0.5 km) composite xenolith in the southern part of the

Marshall Gulch pluton (Fig. 1.3). It yielded a simple population of equant, light pink to clear zircon. Some of the grains are euhedral with well-defined faces, but most are subhedral to anhedral fragments. Although this sample was collected to determine the timing of metamorphism accompanying deformation in the

Marshall Gulch domain, cathodoluminescence (CL) imaging reveals that the zircons are internally characterized by concentric zonation that is generally interpreted to reflect igneous growth (Fig. 1.8C inset). Three zircon fractions (Z1,

32 Z3, and Z5) define a line with intercepts of 1693±2 Ma and 51±351 Ma (Fig.

1.8C). A fourth fraction (Z2) plots beneath this line and is presumed to contain an inherited component. The upper intercept is interpreted as the age of crystallization of the tonalite, and the lower intercept age likely reflects Pb loss related to Laramide tectonism or Rio Grande rifting.

PEGMATITE DIKE (J01-MG3). This sample was collected from one of the thick

(10 – 15 meter), NW-striking, subvertical pegmatite dikes cutting across the southern part of the Marshall Gulch pluton (Figs. 1.3 & 1.7). It yielded a simple population of pink to colorless, equant zircon, nearly half of which are euhedral and display prismatic faces that are typical of igneous growth. The other half of the population consists of subhedral grains with prismatic faces and anhedral fragments interpreted to represent remnants of larger grains. Three zircon fractions (Z1, Z2, and Z4) define a line with intercepts of 1407±6 Ma and 20±45

Ma (Fig. 1.8D). One zircon fraction (Z3) plots just above the regressed line and is not included in the calculation. The upper intercept is interpreted to represent the time of crystallization, and the lower intercept is attributed to recent Pb loss, likely related to Laramide tectonism or Rio Grande rifting.

33 Crestone area

CRESTONE QUARTZ MONZONITE (J01-RA60). A sample of medium-grained, weakly foliated quartz monzonite was collected 0.5 km from the mouth of Burnt Gulch east of Crestone (Fig. 1.5). It yielded a simple population of colorless to light tan, equant, prismatic, euhedral to subhedral zircon as typical of an igneous origin. Four zircon fractions (Z1 – Z4) define a line with intercepts of 1682±3 Ma and 23±20 Ma (Fig. 1.9A). The upper intercept is interpreted to represent the age of igneous crystallization, and the lower intercept likely represents recent Pb loss, probably related to heating and/or fluid flow during Rio Grande rifting. This sample also yielded abundant pale yellow, angular titanite fragments. Three titanite fractions (T1, T4, and T5) define a line with intercepts of 1420±4 Ma and 324±170 Ma (Fig. 1.9A). A fourth fraction (T2) does not fall on this line but, instead, overlaps concordia with a 207Pb/206Pb age of 1489 Ma. Both of these ages are significantly younger than the age of igneous crystallization. In thin section, quartz displays evidence for extensive grain boundary migration with minor rotational recrystallization, and feldspar grains are dominantly cracked with some fine-grained recrystallization along the fractures. These microtextures along with metamorphic mineral assemblages in gneissic host rocks indicate that temperatures did not exceed upper greenschist facies conditions following crystallization of the pluton, and such temperatures are well beneath accepted U-

Pb closure temperatures for titanite (~700°C; Pidgeon et al., 1996; Verts et al., 1996). The upper intercept of the younger titanite population is interpreted to represent metamorphic recrystallization synchronous with development of the

34 solid-state biotite and quartz foliation in the monzogranite. The lower intercept represents more recent Pb loss, likely related to Ancestral Rockies tectonism. The older titanite fraction (T2) might represent a separate, poorly preserved metamorphic recrystallization event at ca. 1489 Ma, but this age falls within a well-documented gap in magmatic and metamorphic ages between ca. 1.63 – 1.47 Ga across the southern Rocky Mountains (Reed et al., 1993).

AMPHIBOLITE (J02-C1). A sample of fine-grained amphibolite was collected from the southern contact aureole of the Crestone stock along South Crestone

Creek 0.25 km from the Willow Creek U.S. Forest Service trailhead, and it was collected to constrain the age of metamorphism in host rocks surrounding the

Crestone quartz monzonite. It yielded a population of colorless zircon with diverse morphologies that are generally equant but include euhedral with prismatic faces, subhedral with and without prismatic faces, blocky and angular anhedral, and rounded anhedral grains. Cathodoluminescence (CL) imaging of approximately 25 grains revealed an equally diverse array of internal structures ranging from well-developed, concentric zonation to irregular, patchy zonation

(Fig. 1.9B insets). Seven zircon fractions representing various morphologies and internal structures were analyzed and suggest that there is no obvious correlation between age, morphology, and internal structure in this sample. Instead, the fractions plot within an envelope of ages defined by an older reference line with intercepts of 1760 Ma and 350 Ma and a younger reference line with intercepts of

35

Figure 1.9. U-Pb concordia diagrams for samples from the Crestone area. Ages are defined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

36 1425 Ma and 0 Ma (Fig. 1.9B). Two fractions (Z3 and Z7) plot along the older line, and the upper intercept is consistent with published ages for basement rocks in neighboring areas (1770 – 1750 Ma, Gunnison-Salida volcanic-plutonic terrane of Bickford and Boardman, 1984). The lower intercept is broadly consistent with the timing of Ancestral Rockies orogenesis in southern Colorado (Kluth and

Coney, 1981; Kluth, 1988). One fraction (Z4) plots along the younger reference line, and the upper intercept is consistent with titanite ages in the adjacent quartz monzonite that are interpreted to represent metamorphic recrystallization during solid-state deformation. The remaining fractions plot between these reference lines and have 207Pb/206Pb ages that range from 1647 – 1509 Ma (Table 1.1).

Although some of the ages suggested by zircon fractions from this sample provide promising correlations, further analyses coupled with careful CL characterization would be required to precisely determine and interpret these ages. These new data are interpreted to represent complex U-Pb systematics likely resulting from multiple periods of zircon growth and varying degrees of recrystallization, regrowth, and/or Pb loss during Paleoproterozoic and Mesoproterozoic magmatism and metamorphism.

Music Pass area

MUSIC PASS QUARTZ MONZONITE (J01-MP1). A sample of foliated, K-feldspar megacrystic quartz monzonite was collected along the ridgeline 0.6 km southeast

37 of Music Pass on the eastern side of the Sangre de Cristo Mountains. It was sampled to determine the minimum age of quartzite deposition because it locally intrudes quartzite and was interpreted by previous workers to be Paleoproterozoic in age (Tweto, 1979; Johnson et al., 1987). This sample yielded a simple population of colorless to light pink, equant, euhedral to subhedral, prismatic zircon interpreted to be igneous in origin. Four fractions (Z1 – Z4) define a line with intercepts of 1434±2 Ma and 0±92 Ma (Fig. 1.10A). The upper intercept is interpreted to represent the age of crystallization of the quartz monzonite, and the lower intercept is attributed to recent Pb loss, probably related to Laramide tectonism or Rio Grande rifting. This sample also yielded abundant brown to dark brown, angular titanite fragments. Three titanite fractions (T1 – T3) define a line with intercepts of 1412±4 Ma and 136±480 Ma (Fig. 1.10A). Whereas quartz monzonite from the sampled locality contains only one foliation that is interpreted to be magmatic in origin (Smagmatic), thin sections from other foliated parts of the pluton contain clusters of titanite that are observed to occur parallel to the recrystallized biotite foliation (S3). Thus, the upper intercept is interpreted to reflect metamorphic recrystallization of titanite, likely related to localized solid-state deformation of the pluton.

PEGMATITE DIKE (J01-MP2). A thin (20 – 30 cm) pegmatite dike was sampled from the same outcrop as the Music Pass quartz monzonite (J01-MP1). The dike sharply cuts the biotite and K-feldspar fabric present in the quartz monzonite, and

38

Figure 1.10. U-Pb concordia diagrams for samples from the Music Pass area. Ages are defined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

39 it was sampled to provide a minimum age on the development of this foliation.

The sample yielded a simple population of colorless to light pink, equant to slightly elongate (1.5:1), euhedral, prismatic zircon that is consistent with an igneous origin. A natural regression of four zircon fractions analyzed (Z1 – Z4) yields a line with a near-zero probability of fit. Two fractions (Z2 and Z4) have

207Pb/206Pb ages (1434 Ma, Table 1.1) that correspond with the age of the host- rock granite (1434±2 Ma, J01-MP1), and fraction Z4 overlaps concordia at 1434

Ma (Fig. 1.10B). Fractions Z1 and Z3 yielded 207Pb/206Pb ages of 1439 and 1440, respectively, and are interpreted to represent inherited zircon. Although these results did not permit precise determination of the age of the pegmatite dike, the two younger fractions suggest that it is not detectably younger than the Music

Pass pluton.

PROTEROZOIC TECTONIC HISTORY OF THE SANGRE DE CRISTO MOUNTAINS

New and existing geochronology and structural data described above provide precise constraints on the age of Proterozoic magmatism, deformation, and metamorphism throughout the Sangre de Cristo Mountains. The Paleoproterozoic tectonic history of the range is characterized by four major events: 1) early intermediate and mafic magmatism accompanied by penetrative deformation and metamorphism, 2) granitic (locally bimodal) magmatism, 3) quartzite deposition, and 4) localized NW – SE shortening and dextral shear

40

Figure 1.11. Schematic block diagrams illustrating the temporal and spatial relationship between Proterozoic mid-crustal magmatism and deformation in the Sangre de Cristo Mountains.

41 accompanied by metamorphism. During the Mesoproterozoic, the range experienced two pulses of granitic magmatism, coeval mafic diking, and one episode of tectonism. New and existing age constraints for these events are summarized in Table 1.2, and schematic block diagrams shown in Figure 1.11 illustrate the spatial relationships between magmatism and deformation through time across the range. The Proterozoic tectonic history of the range is summarized in the following section along with new geochronology constraining the relative and absolute ages of each event.

Early magmatism, deformation, and metamorphism (I1/D1/M1)

Combined with existing geochronology from the southern part of the range, new results described above suggest that basement gneisses throughout the central and southern Sangre de Cristo Mountains formed at approximately the same time. The oldest zircon fractions in amphibolite from the Crestone area yielded ages of ca. 1760 Ma (Fig. 1.9A), and these zircon grains are internally characterized by concentric zonation interpreted to form during igneous crystallization. Thus, this age is interpreted to represent crystallization of the protolith in the central part of the range. Compositionally similar gneisses are intruded to the south by the 1750 – 1730 Ma Blanca Peak suite of plutonic rocks (Fig. 1.2), and deposition and/or crystallization of these rocks must predate the oldest intrusive unit (1749±4 Ma Quartz diorite; Sabin, 1994). Metarhyolite and other assorted metavolcanic rocks comprising basement gneisses to the north yielded U-Pb zircon ages of ca. 1728 – 1670 Ma (Fig. 1.2; Bickford et al., 1989b),

42 suggesting that basement gneisses exposed along the length of the Sangre de Cristo Mountains might represent at least two different age groups (i.e., pre-1750 Ma and post-1730 Ma). This interpretation might, in part, explain the general compositional trend from felsic to mafic from N – S, but a contact between the two assemblages has not been recognized.

The Blanca Peak suite of intrusive rocks (I1) provides approximate age constraints on the earliest phase of deformation and metamorphism (D1 and M1) that affected the Sangre de Cristo Mountains. Published geologic mapping (Johnson and Bruce, 1991; Bruce and Johnson; 1991; Johnson et al., 1987) and reconnaissance mapping by this author in the Blanca Peak area (Figure 1.2) indicate that the suite of 1750 – 1730 Ma intrusive rocks cuts the subvertical,

NW-striking fabric (S1) that characterizes exposures throughout much of the range. However, S1 is also locally deflected around the pluton margins. The Blanca Peak intrusive rocks locally contain a foliation that it is subparallel with foliation patterns in the host-rock gneisses. These relationships suggest that D1 involved subhorizontal, NE – SW shortening that was broadly synchronous with emplacement of the Blanca Peak intrusive suite between 1750 – 1730 Ma. The penetrative gneissic foliation developed during D1 suggests that deformation was accompanied by high-temperature metamorphism (M1) throughout the range during the same time. Based on the close association between deformation, metamorphism, and intermediate to mafic magmatism, D1 and M1 are interpreted to have occurred in an arc environment, perhaps during NE – SW convergence.

43 Granitic magmatism (I2)

The Marshall Gulch monzogranite and Crestone quartz monzonite were emplaced at 1695±2 Ma and 1682±3 Ma, respectively (Figs. 1.8A & 1.9A), and represent a second generation of Proterozoic plutonism (I2) in the Sangre de

Cristo Mountains. Both granitoids locally cut the gneissic fabric (S1), thus requiring that it developed prior to ca. 1692 Ma. Foliated tonalite exposed within the southern part of the Marshall Gulch pluton (Fig. 1.3) is interpreted to have been emplaced at 1693±2 Ma (Fig. 1.8C). If these units are time correlative, magmatism was locally bimodal in character. The Marshall Gulch monzogranite was emplaced as tabular, elongate body parallel with the NW-striking, gneissic host-rock fabric (S1; Fig. 1.3), and the pluton also contains a well-developed magmatic fabric defined by elongate K-feldspar megacrysts that is parallel with

S1. This relationship suggests that magmatic flow occurred parallel to the preexisting gneissic foliation, especially near the margins of the pluton. The Crestone quartz monzonite was similarly emplaced as an elongate body parallel with S1, but it does not contain a strong magmatic fabric and cuts more sharply across the host-rock foliation.

Quartzite deposition

Quartzite exposures are limited in the Sangre de Cristo Mountains.

Underlying gneisses contain evidence for deformation (F1 isoclinal folds) that was not recognized in the quartzite, thus requiring deposition after ca. 1730 Ma. The interpretation that basement also includes Paleoproterozoic intrusive rocks would

44 additionally require deposition after emplacement of the 1682±2 Ma Crestone stock. However, the quartzite is not intruded by any of these older plutons and is instead cut only by the 1434±2 Ma Music Pass pluton. These new constraints from the central Sangre de Cristo Mountains require a range of possible maximum ages of ca. 1730 – 1682 Ma and an absolute minimum depositional age of 1434 Ma. An older (i.e., Paleoproterozoic) minimum quartzite depositional age is inferred based on cross-cutting relationships and structural arguments from the Sangre de Cristo Mountains. The observation that quartzite is intruded by the Music Pass pluton not only requires that deposition predated magmatism, but it also requires that the quartzite was buried to granite emplacement depths (10 – 15 km) prior to ca. 1434 Ma. D2 deformation involved NW – SE shortening in the central part of the range at ca. 1637 Ma (see below), and this age is within the time window for quartzite deposition that is described above. Quartzite deformation and burial is inferred to have occurred during this time, thus requiring that quartzite was deposited during the time window between 1682 – 1637 Ma.

Localized NW – SE shortening, dextral shear, and metamorphism (D2/M2)

After crystallization, deformation (D2) of the Marshall Gulch monzogranite and tonalite involved development of a thick (1 km), subvertical zone of NE-striking fabrics (S2) accompanied by metamorphism at upper greenschist- to lower amphibolite-facies conditions (M2). Metamorphism and

45 deformation are manifested only along the southern margin of the intrusion and involved the development of a subvertical, NE-striking foliation (S2) in the granite. Development of S2 was accompanied by tight, upright folding of both granitic dikes and host-rock gneisses and by local dextral-oblique shearing along a shallow, SW-plunging mineral lineation (L2; Fig. 1.3C). The style of D2 deformation and orientation of S2 fabrics suggest that deformation involved subhorizontal, NW – SE shortening with a component of oblique, dextral strike- slip displacement. Metamorphic titanite clustered parallel to S2 in the deformed granite constrains the age of metamorphism (M2) accompanying deformation to 1637±6 Ma (Fig. 1.8B).

Granitic magmatism and mafic diking (I3)

The coarse-grained Music Pass quartz monzonite represents the third episode of plutonism (I3) in the Sangre de Cristo Mountains. New results indicate that the pluton was emplaced at 1434±2 Ma (Fig. 1.10A) as a tabular, elongate body parallel with gneissic foliation (S1) in the host rocks. A WNW-striking magmatic fabric defined by aligned K-feldspar megacrysts parallels S1 (Plate 1). This relationship suggests that magmatic flow occurred parallel to the host rock fabric. The pluton is cut by a thick (30 m), subvertical mafic dike that strikes NW and contains no internal fabric. Emplacement of the dike is interpreted to have been contemporaneous with granitic magmatism, but an attempt to date it was unsuccessful.

46 Localized NW – SE shortening and metamorphism (D3/M3)

After crystallization of the Music Pass pluton, numerous, discrete zones of solid-state, and locally mylonitic, fabrics were developed during a third deformation event (D3). Subvertical zones of solid-state deformation are up to 1 m thick and strike NE – SW, and the boundary between deformed and undeformed granite is gradational over tens of centimeters. Local asymmetric fabrics include composite foliation (C-S fabric) and pegmatite dikes are offset by up to a few meters, but the presence of both dextral and sinistral kinematic indicators suggest that D3 involved local conjugate shearing during subhorizontal, NW-directed bulk shortening. The southern margin of the 1682±2 Ma Crestone quartz monzonite contains similar, coaxial, solid-state fabrics, and host-rock amphibolite and gneiss were broadly folded (100 meter wavelength) against the pluton, producing moderately NE-plunging axes. The orientation of fabrics and folds suggest that D3 involved subhorizontal, NW – SE shortening. Timing constraints from metamorphic titanite recrystallized parallel with fabrics in the two granitic bodies constrain D3 and M3 to have occurred between 1420 – 1412 Ma.

Pegmatite diking (I4)

The youngest recognized Proterozoic intrusive phase in the Sangre de Cristo Mountains is the swarm of thick (10 – 15 m) pegmatite dikes that cuts the southern part of the Marshall Gulch pluton. These dikes are subvertical, strike NW – SE, and sharply cut all of the fabrics present in the granite and host

47 gneisses. New results indicate that they were emplaced at 1407±6 Ma (Fig. 1.8D). A map-scale intrusion associated with these dikes has not been identified. The parental pluton is interpreted to occur beneath the current level of exposure in the Marshall Gulch area (Fig. 1.11).

REGIONAL (SW US) CORRELATIONS

The tectonic history described above corresponds well with regional evidence constraining the tectonic evolution of the southwestern U.S. during the Proterozoic. In general, basement gneisses of the Sangre de Cristo Mountains are texturally, compositionally, and geographically correlated with two successions of metavolcanic and metasedimentary rocks exposed along a belt extending across southwestern Colorado. Ca. 1750 Ma and older gneisses in the central and southern Sangre de Cristo Mountains are correlated with the Dubois succession, a bimodal suite of metavolcanic rocks that formed between 1770 – 1760 Ma in an island-arc setting (Condie and Nuter, 1981; Bickford and Boardman, 1984; Knoper and Condie, 1988). Younger metavolcanic rocks from the northern part of the range are correlated with the Cochetopa succession. This basement assemblage includes felsic metavolcanic rocks and volcaniclastic sediments with interlayered amphibolites that were formed between 1745 – 1730 Ma in a continental-margin arc setting (Bickford and Boardman, 1984; Knoper and Condie, 1988).

The timing and orientation of D1 and M1 in the Sangre de Cristo Mountains are consistent with other early deformational and metamorphic events

48 that have been documented across southern Colorado. The Iris syncline, south of

Gunnison, Colorado, is a kilometer-scale F1 fold characterized by a NW-trending axial surface and a steep, SE-plunging fold axis (Afifi, 1981). Average S1 surrounding the fold is subvertical and strikes NW – SE with an associated mineral lineation that plunges moderately to steeply NW. Formation of the Iris syncline was synchronous with high-temperature metamorphism (M1; Afifi, 1981) and is constrained by cross-cutting igneous bodies to have formed between 1740 – 1725 Ma (Bickford and Boardman, 1984). Jessup et al. (2005) documented early subvertical, penetrative NW-striking fabrics (S1) and NW-trending folds (F1) in the Black Canyon of the Gunnison. Metamorphic zircon obtained from an amphibolite within the Black Canyon Succession of metasedimentary and metavolcanic gneisses yielded an age of ca. 1742 Ma and is interpreted to represent metamorphism (M1) accompanying development of these early fabrics (Jessup, 2003). These various events suggest that an important regional episode of NW – SE shortening and high-temperature metamorphism occurred between ca. 1750 – 1725 Ma that preceded the development of the pervasive, NE-striking tectonic grain during subsequent orogenic episodes (Karlstrom and Bowring, 1988, 1993; Karlstrom and Humphreys, 1998). Granites (and locally tonalite) intruded into the central Sangre de Cristo Mountains between 1695 – 1682 Ma are correlated with a voluminous suite of syn- to post-Yavapai granitoids emplaced between 1705 – 1666 Ma throughout southern Colorado (Arkansas River Gorge suite of Anderson and Cullers, 1999; Bickford et al., 1989a). The oldest of these plutons contain the strongest

49 penetrative fabrics of the intrusive suite and were emplaced during the waning stages of Yavapai-aged deformation and metamorphism. Younger granites of this inferred late- to post-orogenic suite are either weakly deformed or undeformed, and their geochemistry has been interpreted to reflect an increasing crustal component through time (Anderson and Cullers, 1999). Whereas observations and inferences described above suggest that quartzite deposition occurred after ca. 1695 – 1682 Ma granitic magmatism in the Sangre de Cristo Mountains, regional constraints suggest that quartzite deposition likely coincided with post-Yavapai granitic magmatism. New U-Pb zircon ages from basement granitoids underlying quartzite along Blue Ridge, Colorado, require that quartzite was deposited on exhumed, mid-crustal rocks after ca. 1700 Ma (see Chapter 2 of this study). Published ages (U-Pb zircon) from rhyolite underlying quartzites in New Mexico and Arizona indicate that the onset of quartzite sedimentation closely followed the Yavapai Orogeny at ca. 1.70 Ga (Bauer and Williams, 1989; Cox et al., 2002b). Absolute minimum age constraints across the region typically require that quartzites were deposited prior to Mesoproterozoic (ca. 1.4 Ga) granitic magmatism, thus defining a ca. 250 m.y. time window for deposition. However, based on regional structural arguments and the style of deformation in quartzites throughout the southwestern U.S., deformation and burial of quartzites is inferred to have occurred during the 1.66 – 1.63 Mazatzal Orogeny (Karlstrom and Bowring, 1988; Williams, 1991; see Chapter 2 of this study for discussion). These absolute and inferred regional constraints define a relatively narrow time window of ca. 50 m.y. for quartzite

50 deposition between the Yavapai and Mazatzal orogenies and agree well with new constraints described above from the Sangre de Cristo Mountains. However, regional evidence for the onset of quartzite sedimentation at ca. 1700 Ma requires that deposition was essentially contemporaneous with granitic (and locally bimodal) magmatism at deeper crustal levels within the Sangre de Cristo Mountains. Furthermore, post-Yavapai granitic magmatism was essentially continuous between ca. 1705 – 1663 Ma across southern Colorado, thus coinciding with the entire time window during which quartzite was likely deposited.

Localized NW – SE shortening (D2) and metamorphism (M2) at 1637±6 Ma in the Sangre de Cristo Mountains is correlated with similar-aged deformation and metamorphism documented in the Homestake shear zone of central Colorado. Shaw et al. (2001) reported a monazite age of 1637±13 Ma from rocks with NE- striking, subvertical fabrics within the shear zone. This monazite and another dated at 1658±5 Ma are interpreted to have grown broadly synchronous with a second phase of deformation (D2), locally involving the development of NE- trending, subvertical foliation domains during deformation dominated by NW – SE contraction and vertical extension (Shaw et al., 2001). Further to the south in New Mexico, N-directed crustal shortening occurred across a series of NE- striking structures after deposition of 1664±3 Ma supracrustal rocks and prior to 1654±1 Ma post-kinematic plutonism (Bauer and Williams, 1994). These observations and regional correlations suggest that a major episode of NW – SE crustal shortening occurred throughout the region between 1660 – 1630 Ma, and

51 this second deformational event is interpreted to represent the Mazatzal Orogeny in southern Colorado. The 1434±2 Ma Music Pass pluton is correlated with a widespread suite of coarse-grained, A-type granites emplaced across the southwestern U.S. between 1440 – 1430 Ma (Reed et al., 1993). Granite plutons emplaced during this time across the region are commonly deformed (e.g., Oak Creek pluton; Bickford et al., 1989a), but local observations indicate that neither the Music Pass pluton nor its host rocks were deformed during emplacement. Solid-state deformation (D3) and amphibolite-facies metamorphism (M3) postdated emplacement, occurring locally between 1420 – 1412 Ma and involving subhorizontal, NW – SE shortening. Deformation during this specific time frame is not widely recognized across the region but appears to have been followed by a second episode of granitic magmatism. NW-striking pegmatite dikes emplaced in the Sangre de Cristo Mountains at 1407±6 Ma are correlated with a swarm of W- to WNW- striking pegmatite dikes was emplaced in the Black Canyon of the Gunnison at 1413±2 Ma (Jessup, 2003). Dike emplacement to the west occurred during the waning stages of dextral, tranpressive deformation and was accompanied by amphibolite-facies metamorphism recorded by titanite growth and/or recrystallization in the 1434±2 Ma Vernal Mesa monzonite (Jessup et al., 2005).

REGIONAL (SW US) TECTONIC IMPLICATIONS

New constraints on the Proterozoic tectonic history of the Sangre de Cristo Mountains, interpreted within a regional tectonic framework, have important

52 implications for processes governing the tectonic evolution of southern Laurentia.

Whereas early (1750 – 1720 Ma), penetrative deformation (D1) affected the entire range and developed fabrics that exerted structural control during subsequent

Proterozoic to Phanerozoic tectonism and magmatism, younger deformation (D2 –

D3) was both temporally and spatially localized throughout the range. These contrasting styles of structural behavior suggest that the dominant architecture of the juvenile crust was formed early in its history prior to final northward accretion along the southern margin of Laurentia. Early, subvertical NW – SE fabrics preserved across large regions of southern Colorado and the southwestern U.S. suggest that crustal growth was accompanied by widespread, NE – SW shortening, perhaps related to subduction along arcuate, NW-striking margins (Jessup et al., 2005) and/or collisional orogenesis outboard (south) of the Laurentian margin (Duebendorfer et al., 2001). Although the newly formed crust was subsequently modified by magmatism and tectonism throughout the Proterozoic, the Sangre de Cristo Mountains behaved as a coherent tectonic block and record strongly localized, albeit temporally and kinematically compatible, responses to these events.

The interpreted temporal overlap between quartzite deposition and granitic magmatism would require that sedimentation occurred in a tectonically active environment following the Yavapai Orogeny. Furthermore, there is a distinct spatial and temporal overlap between surface quartzite sedimentation and nearly continuous post-Yavapai granitic (locally bimodal) magmatism throughout southern Colorado. Granitic plutons emplaced between ca. 1695 – 1666 Ma were

53 essentially undeformed during emplacement and only locally contain younger, solid-state fabrics. Anderson and Cullers (1999) argued that the geochemical characteristics of these plutons are consistent with within-plate, post-orogenic magmatism involving an increasing crustal component through time. Age constraints on quartzite deposition and regional structural arguments for the age of quartzite deformation suggest that regional sedimentation occurred within a ca. 50 m.y. time window following the Yavapai Orogeny (Cox et al., 2002b; Chapter 2 of this study). Locally thick (1 – 2 km) quartzite sequences require that relatively deep depositional basins must have been developed to accommodate the influx of sediment. The observation that quartzite is locally deposited on Yavapai-aged granitoids requires tens of kilometers of exhumation of mid-crustal rocks in the 5 – 10 m.y. preceding deposition. These relationships all suggest that local, if not regional, crustal extension might have occurred during this time, perhaps accounting for voluminous magmatism at deeper crustal levels and widespread sedimentation at the surface of the crust. Although post-Yavapai extension might have facilitated the development of depositional basins at the surface of the crust, sediments deposited in this type of tectonically active environment would not be expected to display the high degrees of compositional maturity that characterize Proterozoic quartzites across the region. Thus, various workers have suggested that unique environmental influences that would enhance chemical alteration are required during the Proterozoic to produce thick sequences of syn-orogenic quartzites (e.g., Medaris et al., 2003). If quartzites were deformed during D2 and buried to granite

54 emplacement depths, one might expect D2 fabrics to be more widespread across the Sangre de Cristo Mountains. However, documented D2 deformation is only locally developed along the southern margin of the Marshall Gulch pluton, suggesting that there might be unrecognized Paleoproterozoic structures (i.e., shear zones) within the range that accommodated larger amounts of displacement during D2. Alternatively, Mazatzal-aged deformation preferentially affected structurally higher crustal levels in southern Colorado (Shaw and Karlstrom, 1999). The latter suggestion is consistent with regional evidence for fold-and- thrust style deformation primarily affecting shallower crustal levels elsewhere across the southwestern U.S. during the Mazatzal Orogeny (e.g., Doe and Karlstrom, 1991). Newly recognized Mesoproterozoic (ca. 1.4 Ga) granitic and mafic magmatism and deformation in the Sangre de Cristo Mountains provides new insights regarding the timing and style of ca. 1.4 Ga deformation. Two pulses of magmatism occurred at 1434±2 Ma and 1407±6 Ma, respectively, and intrusive bodies were emplaced along the pre-existing, NW-striking gneissic foliation or as a suite of NW-striking, subvertical dikes. Solid-state deformational fabrics (S3) contained within the 1682±2 Ma Crestone stock and the 1434±2 Ma Music Pass pluton suggest that at least one episode of subhorizontal, NW – SE shortening occurred between 1420 – 1412 Ma. The timing of metamorphism and shortening in the Music Pass pluton (1412±4 Ma, metamorphic titanite parallel with S3) overlaps within error with the emplacement age of NW-striking pegmatite dikes 10 km to the north (1407±6 Ma). These various magmatic and structural elements

55 are all broadly compatible with NW – SE regional shortening at ca. 1.4 Ga (Nyman et al., 1994) but reveal temporal, and perhaps spatial, variations in the relative magnitudes of principal stresses locally. The orientation of pre-existing, host-rock foliation at high angles to compressional stresses likely facilitated the emplacement and controlled the geometry of granites (Nyman and Karlstrom, 1997). The inhomogeneous nature of deformation and subvertical fabric orientations are consistent with other regional examples of ca. 1.4 Ga deformation but contrast sharply with evidence for shallowly-dipping, penetrative fabrics in the southern Wet Mountains 30 km to the east (see Chapter 3 of this study). These penetrative fabrics were developed between 1436 – 1360 Ma and are interpreted to represent subhorizontal flow at relatively deep crustal levels. Ca. 1.4 Ga deformation in the Sangre de Cristo Mountains is interpreted to represent a shallower mid-crustal response to regional shortening, and the transient nature of magmatism and deformation locally might represent discrete events occurring along a distal convergent margin.

SUMMARY AND CONCLUSIONS

Exposures across the Sangre de Cristo Mountains variably record magmatism, deformation, and metamorphism related to the southward growth of Laurentia during the Proterozoic. Two major pulses of plutonism, broadly syn- orogenic quartzite sedimentation, and two phases of deformation and metamorphism characterize the Paleoproterozoic history of the range and provide new constraints on the timing, style, and kinematics of deformation related to

56 accretionary orogenesis. An early episode of penetrative deformation and metamorphism was broadly synchronous with calc-alkaline arc magmatism between 1750 – 1720 Ma. Subvertical, NW – SE striking fabrics developed during early deformation (D1) are interpreted to represent early growth and collision of arc terranes and/or crustal provinces (i.e., Mojave and Yavapai Provinces) along NW-striking subduction boundaries outboard (south) of the active Laurentian accretionary margin (Duebendorfer et al., 2001; Jessup et al., 2005). Two newly dated plutons from the central part of the range are correlated with a regional suite of 1705 – 1666 Ma late-to post-Yavapai granites, and quartzite deposition at the surface of the crust was broadly contemporaneous during this time. The spatial and temporal association between quartzite deposition and post-orogenic granitic (and locally bimodal) magmatism suggests that widespread crustal extension might have followed the Yavapai orogenic peak. Subhorizontal, NW – SE shortening and dextral shear at 1637±6 Ma is interpreted to represent Mazatzal-aged orogenesis locally. The Mesoproterozoic history of the Sangre de Cristo Mountains involved two episodes of granitic magmatism at 1434±2 and 1407±6 Ma accompanied by localized mafic diking. Strongly localized, solid-state to locally mylonitic deformation (D3) occurred between 1420 – 1412 Ma. The geometry of ca. 1.4 Ga intrusions and orientation of fabrics developed during D3 suggest that ca. 1.4 Ga granitic magmatism occurred during subhorizontal, NW – SE directed shortening. Furthermore, the inhomogeneous nature of deformation is consistent with regional, coeval styles of deformation interpreted to represent a relatively shallow

57 mid-crustal response to distal convergent tectonism. Through continued field study combined with precise geochronology, it will be possible to further evaluate the regional temporal and kinematic organization of deformation and magmatism during an otherwise protracted ca. 1.4 Ga tectonothermal event. Increased temporal resolution will additionally permit more rigorous evaluation of tectonic models for widespread Mesoproterozoic granitic magmatism throughout southern Laurentia.

58

59

60

61 Table 1.2. Summary of Proterozoic tectonic events in the Sangre de Cristo Mountains, Colorado

Age Event Interpretation Reference Deposition and/or crystallization of basement Bickford and wall rock to Blanca Peak intrusive suite; Boardman >1750 Ma correlative with Dubois succession to west (1984) (1770 – 1760 Ma)

Emplacement of calc-alkaline suite of plutons Sabin (1994) 1750 – 1730 in Blanca Peak area coeval with metavolcanic Bickford and I Ma 1 rocks of Cochetopa succession in northern Boardman part of range and areas to west (1984)

Early deformation and metamorphism This study, broadly synchronous with plutonism in Map patterns of 1750 – 1730 D /M southern part of range; isoclinal folding and Johnson et al. Ma 1 1 development of penetrative, NW-SE striking (1987) subvertical foliation

Granitic magmatism, bimodal monzogranite + 1695 – 1682 I tonalite magmatism in central part of range w/ This study Ma 2 NW-SE subvertical magmatic foliation

Approximate regional and inferred local See Chapter 2 1700 – 1650 constraints on quartzite deposition and references Ma therein

Development of NE-SW subvertical foliation and recrystallization of metamorphic titanite across southern margin of Marshall Gulch This study D /M 1637±6 Ma 2 2 pluton; NW-SE directed shortening w/ minor local component of dextral shear; inferred deformation of quartzite

A-type granitic magmatism w/ NW-SE, 1434±2 Ma I3 subvertical magmatic foliation; NW-striking This study subvertical mafic dike emplaced (?)

Development of NE-SW subvertical foliation and metamorphic recrystallization of titanite 1420 – 1412 in Crestone and Music Pass granite; NW-SE This study D /M Ma 3 3 directed shortening, locally mylonitic w/ minor component of dextral shear in M.P. granite

NW-striking, subvertical pegmatite diking This study I 1407±6 Ma 4 within Marshall Gulch granite

62 Chapter 2. New age (U-Pb and Pb-Pb) constraints on the deposition and deformation of Proterozoic quartzites across the southern Rocky Mountains

ABSTRACT

New field studies and geochronology (U-Pb and Pb-Pb) constrain the age of deposition of Proterozoic quartzites across southern Colorado and northern New Mexico and provide new detrital zircon age information regarding their sedimentary provenance. Quartzite exposed along Blue Ridge, Colorado, was deposited on coarse-grained granitoid basement with ages (U-Pb zircon) of 1705 – 1698 Ma and was folded into a tight, upright NE-trending syncline prior to emplacement of pegmatite dikes at ca. 1436 Ma. Detrital zircon from quartzites are characterized by a single population with a relatively narrow range of

207Pb/206Pb ages between 1.80 – 1.70 Ga. Archean detrital zircon was present in most cases but represented only a small fraction of grains analyzed (<5%). Peak detrital zircon ages among six samples yielded a narrow range of 207Pb/206Pb ages between 1.76 – 1.70 Ga. The youngest detrital zircon analyzed, and, thus, the maximum depositional age of quartzite, was consistently ca. 1700 Ma. These results, combined with published age constraints from across the southwestern U.S., require that quartzite deposition occurred during and after the ca. 1.70 Ga Yavapai Orogeny. Regional structural arguments and the style of quartzite deformation suggest that quartzites were deformed during the ca. 1.65 Ga Mazatzal Orogeny, thus requiring deposition during a narrow (ca. 50 m.y.) time

63 window in a tectonically active environment. The juvenile detrital character of quartzites is consistent with early-cycle, syn-orogenic sedimentation but contrasts sharply with their extreme compositional maturity, perhaps requiring special environmental influences during deposition that enhanced chemical weathering. Similarities among quartzites exposed throughout the southwestern U.S. and along the Laurentian margin suggest that they represent a widespread regional, and perhaps global, episode of sedimentation during the Proterozoic.

INTRODUCTION

One of the more enigmatic aspects of the Proterozoic tectonic evolution of southern Laurentia is the timing and depositional setting of locally-thick (1 – 2 km), supermature quartzite. Quartzites commonly occur in Proterozoic orogens around the world, with some of the best examples exposed across the southwestern U.S. including the Ortega Formation (New Mexico), Mazatzal Formation (Arizona), and Uncompaghre Formation (Colorado). Similar sequences also occur in more localized areas throughout the southwestern U.S. and elsewhere across North America (Baraboo interval, Lake Superior region; Labrador) (Fig. 2.1). These Proterozoic quartzites exhibit sedimentalogical characteristics suggesting that they were deposited along a passive margin (Soegaard and Eriksson, 1985, 1989), but cross-cutting relationships and regional structural arguments suggest that they were deposited during Paleoproterozoic

64

Figure 2.1. Paleoproterozoic quartzites of southwestern Laurentia. Inset shows possible western Rodinian continents and global Proterozoic quartzite occurrences (relative positions all uncertain, modified from Karlstrom et al., 2001; Sears and Price, 2000; Giles and Betts, 2000). Quartzite localities mentioned in text are labeled accordingly.

65 accretionary orogenesis (Williams, 1993; Williams et al., 2003). Precise age constraints on the deposition and deformation of these distinctive lithologic sequences are generally lacking, and existing relative age constraints only require that they were deposited after the 1.7 Ga Yavapai Orogeny but before a widespread magmatic event at ca. 1.4 Ga (e.g., Cox et al., 2002b; Barker, 1969). Furthermore, the limited geographical extent of quartzite exposures complicates the correlation of sequences at both the regional and global scale. Despite these uncertainties, Proterozoic quartzites are interpreted to represent a regionally important episode of sedimentation across the southwestern U.S. Thus, the timing of deposition, the depositional setting, and the provenance of these metasedimentary sequences all have important implications for regional tectonic models for the Proterozoic evolution of southern Laurentia. This study was undertaken to constrain the deposition and deformation of quartzites exposed across southern Colorado and northern New Mexico (Fig. 2.2) and to provide new detrital zircon age information regarding their provenance. Methods included field mapping, conventional ID-TIMS U-Pb geochronology on underlying and cross-cutting igneous bodies exposed at Blue Ridge, Colorado, and detrital zircon geochronology from a suite of quartzites exposed across the region (Fig. 2.2) using laser ablation ICP-MS techniques. New results reveal that quartzite was deposited after 1706±5 Ma and contains detrital zircon populations recording derivation from local, Paleoproterozoic-aged basement with ages as young as ca. 1.70 Ga. Minimum age constraints require that quartzites were both deposited and folded into tight, upright synclines prior to regional granitic

66 magmatism at ca. 1.43 Ga. Based on regional structural arguments, quartzite deposition is inferred to have occurred prior to the ca. 1.65 Ga Mazatzal Orogeny. However, new results described herein do not rule out younger (i.e., post- Mazatzal) deposition.

GEOLOGIC BACKGROUND

The core of Laurentia comprises Archean blocks that were assembled between 2.0 – 1.8 Ga across Proterozoic mobile belts (Hoffman, 1988). Following the assembly of its cratonic core, Laurentia subsequently underwent a period of southward growth from 1.8 – 1.6 Ga during which much of the Precambrian continental lithosphere of the southwestern U.S. was formed. Crustal growth occurred as island arcs, perhaps analogous to the Indonesian region (Jessup et al., 2005), were formed, accreted, and stabilized against the southern margin of the craton. Two regionally recognized orogenic events, the Yavapai Orogeny (1.7 Ga) and the Mazatzal Orogeny (1.65 Ga), punctuated this protracted history of southward growth (Silver, 1965; Karlstrom and Bowring, 1988). Reactivation of the newly-accreted lithosphere occurred in an intraplate setting accompanying a widespread regional pulse of A-type granitic magmatism during the Mesoproterozoic (1.45 – 1.36 Ga; Reed et al., 1993), possibly in response to renewed crustal accretion along a distal southern margin (Nyman et al., 1994). These events were all part of a prolonged (ca. 800 m.y.) episode of crustal growth along southern Laurentia culminating in

67

Figure 2.2. Paleoproterozoic quartzite exposures and inferred regional extent across Colorado and New Mexico. Quartzite localities sampled for geochronology and mountain ranges discussed in the text are labeled accordingly.

68 the Grenville Orogeny and assembly of the supercontinent Rodinia at ca. 1.1 Ga (Karlstrom et al., 2001). During the southward growth of Laurentia, thick (1 – 2 km) sequences of quartz arenite were deposited along the southern margin of the Paleoproterozoic continent. Some of the best examples of these deposits are preserved in the Lake Superior region of the northern U.S. (Baraboo interval) and throughout the southwestern U.S. (Figs. 2.1 & 2.2). Southwestern U.S. quartzite sequences include the Ortega Formation (New Mexico), Mazatzal Formation (Arizona), and Uncompaghre Formation (Colorado), as well as numerous smaller exposures across the region. Quartzites in the metasedimentary sequences are commonly nearly pure (>95% quartz) with minor muscovite, Al-silicates, hematite, zircon, and monazite. Primary sedimentary structures are locally well preserved, including common cross-stratification. Depositional facies are similar from top to bottom and region to region and indicate shallow marine (<10 m water depth) or fluvial environments (Trevena, 1979; Soegaard and Eriksson, 1985). There is a general absence of horizontal or vertical lithofacies variations that would reflect transgressive and regressive cycles. Across much of the southwestern U.S., these quartzites directly overlie thick sequences of voluminous, high-silica rhyolite. The contact between rhyolite and quartzite is generally interlayered to gradational and is commonly marked by a distinctive, Mn-rich contact interval. Throughout southern Colorado, however, rhyolite is notably absent, and quartzites are typically underlain by pebble to cobble conglomerates up to a few meters thick. These basal conglomerates contain a variety of clast compositions that include

69 vein quartz, quartzite, jasper, chert, and, locally, granite (Barker, 1969; Reuss, 1974). Although these various compositions do not always reflect the local makeup of underlying basement assemblages, the uppermost part of the basement is commonly marked by a zone of deep weathering (regolith) interpreted to have developed prior to the onset of quartzite sedimentation. Basement underlying quartzite sequences regionally comprises a very different lithotectonic assemblage characterized by metamorphosed mafic volcanic rocks and volcanogenic marine metasedimentary rocks (Bauer and Williams, 1989). Examples of these older (1.80 – 1.72 Ga; Condie, 1982) sequences include the Moppin Complex in northern New Mexico (Bauer and Williams, 1989) and the Dubois and Cochetopa successions in southern Colorado (Bickford and Boardman, 1984). Basement assemblages commonly contain evidence for multiple episodes of deformation and/or metamorphism that are not recognized in the overlying quartzite. Thus, the contact between basement assemblages and quartzite sequences has been variably interpreted as an unconformable depositional contact or a fault contact. These interpretations obviously have very different implications for constraining the age of quartzite deposition. After deposition, both quartzite sequences and underlying basement assemblages were deformed by folding and thrust imbrication. Across Colorado, quartzite exposures presently occur as tight, upright synclinal “keels” interpreted to be the roots of much larger folds that have been eroded away (Reuss, 1974; Wells et al., 1964). In the Needle Mountains and Tusas Mountains of

70 southwestern Colorado and northern New Mexico, respectively, thick (1 – 2 km) sections of quartzite are folded into tight to open, large-wavelength (kilometer scale) folds consistent with some fold-and-thrust-belt geometries (Harris, 1990; Williams, 1991). During deformation, quartzites were buried to depths of up to 10 – 15 km, and they resided at these mid-crustal depths until at least ca. 1.4 Ga, when they were widely intruded by coarse-grained granitic plutons. Whereas published absolute age constraints (U-Pb zircon) from Arizona and New Mexico indicate that quartzite deposition occurred after the ca. 1.7 Ga Yavapai Orogeny (Cox et al., 2002b; Bauer and Williams, 1989), the minimum age of quartzite deposition is poorly constrained because cross-cutting igneous bodies generally yield ages that are too young (i.e., ca. 1.4 Ga) to narrowly determine the age of deformation.

BLUE RIDGE QUARTZITE, COLORADO

Proterozoic quartzite, schist, and conglomerate are exposed in a NE- trending, tight, upright syncline along Blue Ridge (Figs. 2.2 & 2.3), 10 km NNW of Cañon City, Colorado. The metasedimentary sequence is in contact with coarse-grained, foliated to locally strongly-foliated granodiorite of the Paleoproterozoic Twin Mountain batholith (1705 Ma; Bickford et al., 1989a) and is cut by numerous pegmatite dikes. This area was chosen for geochronologic study to determine the age of quartzite deposition and deformation. The following sections describe the characteristic lithologies present within the

71

Figure 2.3. Generalized geologic map of the Proterozoic quartzite-schist sequence exposed along Blue Ridge, Colorado (from Reuss, 1970, 1974). Location of geochronology samples are indicated by stars. 72

Figure 2.4. Schematic geologic cross sections across Gooseberry Gulch syncline, Blue Ridge, Colorado (from Reuss, 1970, 1974).

73 metasedimentary sequence, the various structural elements present in outcrop, and new U-Pb geochronology results from underlying and cross-cutting granitoids.

Rock descriptions

The Proterozoic metasedimentary sequence exposed along Blue Ridge was mapped and described by Reuss (1974), and the following unit descriptions are summarized largely from his detailed report. Quartzite and schist are exposed in eight NE-trending, subvertical layers ranging in thickness from 15 – 350 m (Fig. 2.3). Quartzite is relatively pure (80 – 95% quartz with 1 – 10% muscovite), and individual quartz grains range in size from 0.5 – 5.0 mm and are strongly recrystallized with interlocking, sutured grain boundaries. Tabular cross-bedding is well preserved with individual bed thicknesses between 0.1 – 0.4 m thick, and paleocurrent analysis corrected for deformation suggests a dominant current direction of NNE to SSW. Oscillation ripple marks are also locally preserved and have an average amplitude and wavelength of 0.5 – 1.5 cm and 2 – 5 cm, respectively. Schist layers are dominated by quartz, biotite, and muscovite but also presently contain a variety of metamorphic minerals including garnet + sillimanite + andalusite + staurolite ± cordierite ± actinolite. Interlayered beds of quartzite and/or quartz pebble conglomerate are common. The quartzite-schist sequence is underlain to the north by a 15 – 25 m thick quartzite pebble conglomerate (Figs. 2.3 & 2.4). The contact between the quartzite and conglomerate ranges from sharp to gradational, and isolated lenses of conglomerate occur within the upper units of the quartzite sequence. Quartzite

74 pebbles are 0.5 – 2.0 cm in size and the surrounding matrix comprises quartz, muscovite, and well-rounded zircon up to 0.1 mm in size. The lower 5 – 10 m of conglomerate also contain up to 20% microcline crystals that are 0.5 – 2.0 cm in size and contain numerous quartz and muscovite inclusions and black tourmaline veinlets. These clasts are interpreted to represent igneous material derived from the underlying granitoid basement. Between the conglomerate and basement granitoid rocks is a thin (5 – 10 m), poorly-exposed zone of phyllitic material that locally contains large granitic clasts (Fig. 2.5). This “regolith” zone is interpreted to represent part of the underlying granodiorite that was sub-aerially exposed and heavily weathered prior to deposition of the conglomerate. Based on the gradational stratigraphic progression from highly-weathered granitoid basement to basal conglomerate to quartzite and schist, the basal conglomerate is interpreted to be in unconformable depositional contact with underlying plutonic basement rocks.

Structural relationships

After deposition, the metasedimentary sequence experienced two major phases of deformation. The first phase (D1) involved development of a bedding- parallel foliation accompanied by growth and/or recrystallization of biotite, muscovite, staurolite, and possibly cordierite (Reuss, 1974). The second phase of deformation (D2) produced the 4 km wavelength Gooseberry Gulch syncline, the dominant structure presently characterizing exposures along Blue Ridge (Figs. 2.3

& 2.4). During D2, the quartzite sequence was folded into a tight-to-isoclinal,

75

Figure 2.5. Schematic illustration of cross-cutting relationships and summary of new U-Pb and detrital zircon geochronology, Blue Ridge, Colorado. For location and map-scale geometry of units, refer to Figure 2.3.

76 upright, NE-trending fold (F2; Fig. 2.4) accompanied by amphibolite-facies metamorphism and local development of a subvertical, axial-planar foliation (S2). It is difficult to determine whether or not basement granitoids were folded with the overlying sequence during D2. The southern limb of the syncline is cut by a ductile shear zone that juxtaposes strongly-deformed, locally-mylonitic granodiorite against steeply-dipping quartzite and schist, omitting nearly 1 km of stratigraphic thickness from the lower part of the metasedimentary sequence (Fig. 2.4; Reuss, 1974). Deformation is most intense within 5 – 10 meters of the granodiorite/quartzite contact and grades into weakly-foliated granodiorite across a distance of 20 – 30 m. The shear zone fabric is subparallel with S2 in the quartzite sequence and is, thus, interpreted to have formed during the late stages of D2 shortening. The folded sequence is also cut by a series of 60 – 180 m left- lateral, en-echelon folds that postdate development of the main syncline. These folds are interpreted to have formed during the late stages of D2 or during a later

D3 (Reuss, 1974).

U-Pb geochronology

A suite of four samples was collected for U-Pb geochronologic analysis to constrain the age of quartzite deposition and deformation at Blue Ridge, Colorado. Samples included foliated granodiorite exposed north of the quartzite sequence, strongly-deformed granodiorite from the southern, sheared contact of the sequence, and two pegmatite dikes that cross-cut the folded sequence and sheared granodiorite (Fig. 2.3 for locations). Quartzite and basal conglomerate

77 were also sampled Pb-Pb geochronology as part of a detrital zircon study described below. U-Pb isotopic data are presented in Table 2.1, and concordia diagrams are presented in Figure 2.6. Zircon fractions were hand picked, examined using a petrographic microscope, characterized by cathodoluminescence, extensively abraded, and then subjected to a final optical re-evaluation before analysis. Complete analytical methods are described in Appendix 1.

FOLIATED GRANODIORITE (J01-BR3). A sample of weakly-foliated, coarse- grained granodiorite was collected 5 km to the north of Blue Ridge and was mapped as part of the Paleoproterozoic Twin Mountain batholith (Fig. 2.3).

Locally, it represents granitoid basement interpreted to unconformably underlie the quartzite sequence and was collected to determine the maximum age for quartzite deposition. The sample yielded a simple population of colorless to light tan, euhedral to subhedral, equant to slightly-elongate prismatic zircon consistent with an igneous origin. Three fractions (Z1 – Z3) define a line with intercepts of

1706+5/-3 Ma and 476±276 Ma (Fig. 2.6A). The upper intercept is interpreted to represent the crystallization age of the granodiorite and is consistent with the published range of ages for similar intrusive phases from elsewhere within the surrounding Twin Mountain and Crampton Mountain batholiths (1706±5 Ma and

1705±8 Ma, respectively; Bickford et al., 1989a). The lower intercept is

78 interpreted to reflect more recent Pb loss likely related to Ancestral Rockies tectonism (Kluth and Coney, 1981; Kluth, 1988).

SHEARED GRANODIORITE (K00-BR-25). Strongly deformed, coarse-grained granodiorite was sampled from the southern contact of the Gooseberry Gulch syncline (Fig. 2.3). Although the nature of the original contact between the quartzite sequence and granodiorite is ambiguous at this location due to shear zone development, this sample was collected to correlate the age of basement granitoids on both sides of the folded quartzite. The sheared granodiorite yielded a single population of colorless to light tan, equant, euhedral to subhedral prismatic grains consistent with an igneous origin. Four fractions analyzed all plot near concordia and have 207Pb/206Pb ages ranging from 1708 – 1693 Ma

(Table 2.1). Three fractions (Z2 – Z4) plot along a reference line with intercepts of 1698 Ma and 0 Ma (Fig. 2.6B). The upper intercept corresponds with the average 207Pb/206Pb age of 1698±4 Ma for the three fractions (Table 2.1) and is interpreted to represent the age of crystallization of the granodiorite. The fourth fraction (Z1) plots below the reference line and is interpreted to contain inherited zircon. The age of this sample overlaps within error the age of granodiorite north of Blue Ridge (sample J01-BR3) and with published ages for the Twin Mountain and Crampton Mountain batholiths, suggesting that the southern contact of the

79

Figure 2.6. U-Pb concordia diagrams for samples from Blue Ridge, Colorado. Ages are determined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

80

81

82 quartzite sequence might have also been an unconformable depositional contact prior to shearing.

CROSS-CUTTING PEGMATITE DIKE (K00-BR-26). This sample was collected from a 1 m thick, white to pink pegmatite dike that cuts across the folded quartzite and sheared granodiorite (Fig. 2.3). This dike is part of a suite of sub-planar intrusions that are commonly thin (<0.5 meters) but, locally, up to a few meters thick. It was sampled to provide a minimum age on D2 deformation and, therefore, deposition of the metasedimentary sequence. The sample yielded a small, relatively simple population of pink to clear, euhedral prismatic zircon consistent with an igneous origin. Two fractions (Z1 and Z2) have 207Pb/206Pb ages of 1686 Ma and 1707 Ma, respectively, and a third fraction (Z3) has a

207Pb/206Pb age of 1436 Ma (Table 2.1; Fig. 2.6C). Although there was no optical evidence for xenocrystic cores or overgrowths, these data suggest that the pegmatite contains inherited zircon. The youngest fraction is interpreted to represent the age of emplacement and crystallization of the pegmatite dike, and the older fractions are interpreted to represent zircon inherited from host-rock granodiorite. Because fraction Z3 does not overlap concordia, its age can only be interpreted as a minimum estimate. However, ca. 1436 Ma is consistent with a widespread regional pulse of granitic magmatism and related pegmatite diking that is well documented both locally and regionally (Reed et al., 1993).

83

CROSS-CUTTING PEGMATITE DIKE (J03-BR4). Because sample K00-BR-26 yielded an age that was so young it did not more narrowly constrain quartzite deposition and deformation, another pegmatite dike was collected following the same rationale as the previous sample. Whereas Proterozoic quartzites across the region are rarely cut by post-depositional intrusions, the folded quartzite-schist sequence exposed along Blue Ridge is cut by numerous pegmatite dikes. Sample

J03-BR4 was collected from a pink to red, thin (30 cm) pegmatite dike that cuts the sheared granodiorite and quartzite fabric but also locally deflects foliation in quartzose schist (Fig. 2.7). The sample yielded a relatively simple population of light pink to clear, euhedral to subhedral prismatic zircon with some elongate (3:1

– 4:1) grains consistent with an igneous origin. Two fractions (Z1 and Z2) have

207Pb/206Pb ages of 1731 Ma and 1699 Ma, respectively, and are interpreted to represent inherited zircon (Fig. 2.6C; Table 2.1). A third fraction (Z3) overlaps concordia with a 207Pb/206Pb age of 1436 Ma (Fig. 2.6C), and this age is interpreted to represent emplacement of the dike. The emplacement age is consistent with the interpreted age of sample K00-BR-26 and with a well- documented regional suite of ca. 1.4 Ga granites and related pegmatite dikes emplaced locally and regionally between 1440 – 1430 Ma (Reed et al., 1993).

84

Figure 2.7. Field photograph of pegmatite dike cutting and deflecting quartzite and schist fabric at Blue Ridge, Colorado (sample J03-BR4, see Fig. 2.3 for location). Hammer handle is approximately 1 m long.

85 QUARTZITE DETRITAL ZIRCON GEOCHRONOLOGY

A suite of Proterozoic quartzites was collected from across southern Colorado and northern New Mexico for a combined Pb-Pb and U-Pb study of detrital zircon to determine the maximum age of quartzite deposition and to characterize the sedimentary provenance of the quartzite sequences. Samples were collected from both thick (1 – 2 km), extensively exposed units (Ortega and Uncompaghre formations) and smaller, more localized occurrences (Cebolla Creek, Blue Ridge, and Phantom Canyon localities) sharing similar characteristics and lithologies. Sample locations are indicated in Figure 2.2. In most cases, representative samples of quartzite were collected from the stratigraphically lowest part of the existing section. Samples of the basal conglomerate and quartzite from 10 meters higher in the stratigraphic section were collected from Blue Ridge, Colorado, to examine possible changes in the detrital population throughout the depositional history of the metasedimentary sequence. All samples were processed according to standard mineral separation procedures described in Appendix 1. Zircon from the least magnetic fraction was hand- picked to include grain populations representing each of the various morphologies, sizes, and colors that were recognized. Approximately 100 zircon grains from each sample were mounted in epoxy discs and lightly polished to expose the grain surfaces. Specific laser ablation (LA) ICP-MS analytical procedures are described in Appendix 2, and data tables are presented in Appendix 3. The estimated uncertainty (i.e., external reproducibility) on each individual analysis is approximately 60 m.y. (3.5%).

86 The results described in the following section are presented as 207Pb/206Pb ages because U/Pb ratios were not rigorously calibrated during LA-ICP-MS analyses. In the absence of corrected U isotopic data, 207Pb/206Pb ages must be interpreted as minimum ages. To determine whether the 207Pb/206Pb ages are accurate and precise within the stated errors, a select group of grains were re- analyzed using isotope dilution thermal ionization mass spectrometry (ID-TIMS) U-Pb methods. The two samples yielding the largest number of young (i.e., <1.70 Ga) grains were chosen, and individual zircon grains with the youngest

207Pb/206Pb ages were removed from the epoxy mounts, lightly air abraded, and analyzed by conventional ID-TIMS methods described in Appendix 1. U-Pb isotopic data are presented in Table 2.1, and U-Pb concordia plots are shown in Figure 2.9.

Results and Interpretations

Zircon morphologies were highly variable among all of the samples and ranged from euhedral and prismatic to well-rounded and lightly-frosted. Approximately 80 grains from each sample were analyzed, and age distribution plots are shown in Figure 2.8. The range of 207Pb/206Pb ages among all samples was 3.38 – 1.52 Ga (Table 2.2), but each sample is primarily characterized by Paleoproterozoic-aged detritus that yielded a relatively narrow range of ages between 1.8 – 1.7 Ga (Fig. 2.8). Archean-aged detritus is present in all samples but one (Blue Ridge basal conglomerate, Fig. 2.8I) but is relatively minor in all cases (up to 5% of grains analyzed). Peak detrital ages similarly define a narrow

87 Table 2.2. Quartzite detrital zircon 207Pb/206Pb age summary

Unweighted Peak age (median) Sample population n Avg. ± % Avg. ± % J01-PC1 All 81 1.72 .13 7.8 1.70 .01 0.76 <2.0 Ga 74 1.71 .05 2.9 1.70 .01 0.74 207Pb/206Pb age range = 2.63 – 1.52 Ma

J03-NM4 All 83 1.75 .12 6.9 1.73 .01 0.75 <2.0 Ga 81 1.74 .06 3.7 1.73 .01 0.76 207Pb/206Pb age range = 2.55 – 1.60 Ma

ORT-N All 81 1.79 .24 13.6 1.73 .01 0.66 <2.0 Ga 75 1.74 .05 2.9 1.73 .01 0.63 207Pb/206Pb age range = 3.38 – 1.65 Ma

J01-BR1 All 78 1.83 .24 13 1.77 .02 1.0 <2.0 Ga 72 1.77 .06 3.4 1.76 .01 0.79 207Pb/206Pb age range = 3.18 – 1.65 Ma

J01-BR2 All 79 1.72 .04 2.5 1.71 .02 0.70 207Pb/206Pb age range = 1.91 – 1.65 Ma

LC-CC-15 All 79 1.83 .24 13 1.77 .02 0.95 <2.0 Ga 74 1.77 .06 3.6 1.76 .01 0.74 207Pb/206Pb age range = 2.95 – 1.65 Ma Peak age represents the statistical median calculated using Isoplot v. 3.09a (Ludwig, 2004). Statistical absolute (±) and percent (%) error are shown at 2σ level. Estimated error on an individual analysis is 3 – 5% (see text and Appendix 2).

88

Figure 2.8. Detrital zircon age distribution plots for Early Proterozoic quartzites from northern New Mexico and southern Colorado. Histograms on left represent the entire detrital zircon population, and histograms on right represent only Paleoproterozoic-aged detrital zircon.

89

Figure 2.8 continued.

90

Figure 2.9. U-Pb concordia diagrams for detrital zircon grains from quartzite samples. Ages are determined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

91 range of ages (1.76 – 1.70 Ga; Table 2.2), and the peak detrital age for each respective sample locality closely matches published U-Pb zircon ages of surrounding basement assemblages. The youngest 207Pb/206Pb ages determined on detrital zircon among all of the quartzite samples were consistently around 1.65 Ga, but a few grains yielded ages as young as 1.52 Ga. However, because these are 207Pb/206Pb ages, they must be treated as minimum ages. Highly-discordant zircon grains can yield a

207Pb/206Pb age, or “apparent” age, that is much younger than a more accurate U- Pb age. This relationship is illustrated by ID-TIMS analysis of the youngest detrital zircon grains from quartzite exposed along Phantom Canyon, Colorado (sample J03-PC1). Three detrital zircon grains yielded relatively young 207Pb/206Pb ages by LA-ICP-MS methods (1.67 – 1.52 Ma), but ID-TIMS analysis reveals that all three grains are variably discordant (up to 42%; Fig. 2.9A). These grains define a mixing line with intercepts of 1701±3 Ma and 2±10 Ma (Fig. 2.9A), and the upper intercept corresponds with the peak 207Pb/206Pb age of the detrital population from the same sample (1.70 Ga; Table 2.2). ID-TIMS analysis of similarly young (1.66 – 1.63 Ga) detrital zircon grains from the Ortega Quartzite (sample ORT-N) demonstrates that these grains are, in fact, nearly concordant but have U-Pb ages that are older by up to 60 – 70 m.y. The range of U-Pb ages (ca. 1725 – 1700 Ma, Table 2.1) is slightly younger than the peak detrital 207Pb/206Pb age of the entire population in this sample (1.73 Ga, Table 2.2). The difference in ages determined by the two techniques reflects the lower

92 precision of the LA-ICP-MS methods used in this study and supports estimated errors of approximately 3.5%, or ca. 60 m.y., for individual analyses. Minor amounts of Archean detrital zircon (up to 5%) are interpreted to reflect sedimentary input from the Wyoming Province to the north but might also reflect reworked sedimentary rocks or unrecognized, Archean-aged crustal fragments that were possibly incorporated into the Paleoproterozoic accretionary orogen (Hill and Bickford, 2001). The bulk of the detritus is Paleoproterozoic in age and defines a relatively narrow range of ages (1.80 – 1.70 Ga). Peak detrital ages vary among the six samples but similarly display a narrow age range (1.76 – 1.70 Ga). The range of detrital zircon ages corresponds with the general age range of juvenile basement assemblages across the region, and the peak detrital ages agree well with the average age of underlying basement assemblages locally. These results are interpreted to reflect the relatively juvenile detrital character of Proterozoic quartzites across southern Colorado and northern New Mexico. Detritus is interpreted to have been derived locally with only restricted input from Archean sources to the north. The transition from a relatively restricted detrital age population in the basal conglomerate (Fig. 2.8I) from Blue Ridge, Colorado, to a broader age range including Archean material from quartzite 10 meters upsection (Fig. 2.8G) suggests that depositional systems likely evolved through time to reach older source terranes. Accordingly, the quartzite from Blue Ridge also yielded one of the oldest peak detrital ages (1.76 Ga, Table 2.2). However, this more evolved system was apparently still dominated by Paleoproterozoic-

93 aged detritus interpreted to represent primary derivation from the 1.80 – 1.70 Ga Yavapai Province. The youngest detrital zircon age and, thus, the maximum age of quartzite deposition, is interpreted to be 1.70 Ga. Although individual grains did yield younger 207Pb/206Pb ages, these data must be treated as minimum ages for reasons described above. Ca. 1.70 Ga represents the youngest peak detrital 207Pb/206Pb age among all samples analyzed (J03-PC1, Table 2.2) and corresponds with the age of the youngest, nearly-concordant zircon analyzed by ID-TIMS methods from the Ortega Quartzite (Fig. 2.9B). Ca. 1.70 Ga also overlaps within estimated error with numerous detrital zircon grains yielding 207Pb/206Pb ages between 1.70 – 1.65 Ga. These results require that quartzites were deposited after the ca. 1.7 Ga Yavapai Orogeny (Karlstrom and Bowring, 1988; 1993), and the presence of 1.70 Ga detrital zircon further suggests that Yavapai-aged basement was exposed at the surface in source areas during deposition. This is consistent with new data described above from Blue Ridge, Colorado, indicating that quartzites were locally deposited on exhumed, ca. 1.7 Ga granitoid basement.

DISCUSSION

Geochronology reported above provides important new constraints on the deposition and deformation of quartzite sequences across southern Colorado and northern New Mexico. These constraints are consistent with quartzites from elsewhere across the southwestern U.S. and North America and suggest that Proterozoic quartzite sequences might represent a widespread regional

94 depositional event. New results are discussed in the following section in the context of regional models for the Proterozoic tectonic evolution of southern Laurentia, and two separate, yet equally permissible, time periods for quartzite deposition are discussed that depend on the age of quartzite deformation.

New constraints on quartzite deposition

New geochronology from exposures along Blue Ridge, Colorado, reveals that granitoid basement underlying quartzite at this locality has U-Pb zircon ages of 1706±5 and 1698±4 Ma (Fig. 2.5). Quartzite is interpreted to be in unconformable depositional contact with the underlying granodiorite based, in part, on a deeply weathered horizon in the upper part of the basement. This “regolith” zone is interpreted to represent sub-aerial exposure of the granodiorite, requiring >10 km of exhumation prior to deposition of the metasedimentary sequence. Deeply weathered granodiorite locally grades into the phyllitic conglomerate defining the base of the quartzite sequence. Thus, ca. 1706 – 1698 Ma is interpreted to represent a robust maximum age for the deposition of quartzite locally, and this interpretation is additionally supported by detrital zircon geochronology from the same location. The basal conglomerate from the sequence exposed along Blue Ridge is characterized by a detrital zircon population with a mean age of ca. 1.71 Ga (Table 2.2), and individual detrital zircon grains consistently yielded 207Pb/206Pb ages as young as 1.70 Ga (within error of LA-ICP-MS methods, see above). Although quartzite sampled from higher (10 m) within the metasedimentary sequence is characterized by a detrital

95 zircon population with a much older peak age (1.76 Ga; Table 2.2), it similarly yielded individual grains with 207Pb/206Pb ages as young as 1.70 Ga. The minimum age of quartzite deposition is newly defined at Blue Ridge by multiple cross-cutting pegmatite dikes that were emplaced at ca. 1436 Ma (Fig. 2.6C). These dikes cut the quartzite sequence and the Gooseberry Gulch syncline, thus requiring deposition of the quartzite and two episodes of deformation (D1 and

D2) prior to dike emplacement at 1436 Ma. The observation that one of the pegmatite dikes (sample J03-BR4) locally deflects foliation in host-rock quartzite and schist (Fig. 2.7) suggests that some degree of deformation, perhaps involving local fabric reactivation, accompanied magmatism at 1436 Ma. However, the dikes are otherwise strongly discordant with respect to the kilometer-scale syncline (F2) that dominates exposures along Blue Ridge. Quartzite sampled from Phantom Canyon, 20 km to the east of Blue Ridge, is characterized by a detrital zircon population with the youngest peak age of all samples analyzed (1.70 Ga, Table 2.2), and ID-TIMS analysis of the three youngest detrital zircon grains from the quartzite yielded a U-Pb age of 1701±3 Ma (Fig. 2.9A). These ages are identical with the age of basement granitoids underlying quartzite along Blue Ridge and confirm ca. 1.70 Ga as the maximum age of quartzite deposition locally. The Ortega Quartzite, sampled in the Tusas Mountains of northern New Mexico, is characterized by a detrital zircon population with a peak age of 1.73 Ga (Table 2.2), and it yielded nearly- concordant, individual zircon grains with ages as young as ca. 1700 Ma. These young ages are consistent with the maximum depositional ages determined from

96 further north in Colorado and suggest that quartzite sequences across a larger region might have been deposited during the same time interval. The overall detrital zircon populations among all quartzites sampled are remarkably similar (Fig. 2.8) and are dominated by Paleoproterozoic-aged detritus. These similarities suggest that the quartzites sampled likely represent correlative metasedimentary sequences, and the variation in the peak detrital age is interpreted to represent differences in the range of basement ages present in the respective source areas. Mazatzal-aged (1.65 – 1.60 Ga) detritus was not recognized in any of the quartzites sampled. Although some detrital zircon grains did yield 207Pb/206Pb ages as young as 1.65 – 1.52 Ga in this study, these ages must be interpreted only as minimum estimates for reasons described above. The youngest U-Pb ages and peak detrital 207Pb/206Pb ages consistently agree at ca. 1.70 Ga, and this age is interpreted to represent the youngest detritus present in the quartzites sampled. Although the absence of Mazatzal-aged detritus could simply represent depositional systems that derived material primarily, if not exclusively, from northern, older (>1.7 Ga) source terranes, it also suggests that quartzites might have been deposited prior to accretion of the Mazatzal Province to the south.

Regional (SW US) constraints on quartzite deposition

New constraints on Proterozoic quartzite deposition agree well with other published constraints from across the southwestern U.S. Zircon from Vadito Group rhyolites underlying quartzites of the Ortega Formation in the Tusas Mountains, New Mexico (Fig. 2.2), yielded U-Pb ages of ca. 1700 Ma (Bauer and

97 Williams, 1989). These rhyolites locally grade into quartz-rich metasedimentary rocks that include trough-bedded quartzites and, thus, are interpreted to represent part of a continuous stratigraphic sequence. In the Mazatzal Group of central Arizona (Fig. 2.1), Cox et al. (2002b) interpreted the onset of quartzite sedimentation to have occurred at 1701±2 Ma based on U-Pb zircon ages from a rhyolite ash-flow tuff interlayered with the lower part of the section. The Mazatzal Group comprises quartzite, conglomerate, and shale deposited unconformably on a volcanic assemblage with ages of 1709±6 Ma and 1700±6 Ma (Cox et al., 2002b; Silver et al., 1986), and the composition of the sedimentary sequence suggests derivation from local orogenic and basement rocks (Cox et al., 2002b). Whereas new maximum age constraints described above only require that deposition occurred after the ca. 1.7 Ga Yavapai Orogeny (Karlstrom and Bowring, 1988, 1993), these U-Pb ages from within continuous stratigraphic sequences suggest that deposition actually closely followed the Yavapai orogenic peak. New minimum age constraints described herein require only that quartzite sequences were deposited and deformed prior to ca. 1436 Ma, and similar constraints are well documented across southern Colorado. In the Needle Mountains (Fig. 2.2), folded quartzite and schist of the Uncompaghre Formation are sharply truncated by the coarse-grained, 1438±20 Ma Eolus Granite (Silver and Barker, 1968; Barker, 1969). In the Sangre de Cristo Mountains (Fig. 2.2), a narrow (100 m) band of cross-bedded quartzite is intruded by the 1434±4 Ma Music Pass pluton (see Chapter 1 of this study), and large (10 – 15 m) quartzite

98 xenoliths are preserved in the coarse-grained quartz monzonite within 0.5 km of the intrusive contact. These relationships not only require that quartzites were deposited prior to ca. 1.43 Ga granitic magmatism, but they additionally require that the sequences were buried to granitic emplacement depths (10 – 15 km) prior to intrusion. Furthermore, ca. 1.43 Ga granites also cut all of the deformation fabrics present in the metasedimentary sequence. Thus, both deposition and deformation of quartzite sequences must have occurred prior magmatism at ca. 1.43 Ga.

Pre-Mazatzal Orogeny (ca. 1.65 Ga) deposition

Absolute age constraints described above define a relatively broad time window for quartzite deposition between ca. 1.70 – 1.43 Ga. However, based on regional structural arguments, a lack of Mazatzal-aged detritus, and deformation styles in the quartzite, burial is inferred to have occurred during the ca. 1.65 Ga Mazatzal Orogeny. Deformation associated with this orogenic episode primarily affected Arizona and New Mexico but also propagated northward into Colorado. The “Mazatzal Front” represents the approximate northern extent of these effects

(Figs. 2.1 & 2.2; Karlstrom and Bowring, 1993; Karlstrom and Daniel, 1993), and all of the quartzites sampled for this study occur south of this deformation front. Williams (1991) described inferred Mazatzal-aged deformation in the Tusas Mountains, northern New Mexico, that involved N-directed km-scale folding and thrusting of the Ortega and Vadito Group quartzite-rhyolite sequence and underlying basement assemblages. Quartzites of the Ortega Formation record a

99 minimum of 50% shortening that was accommodated primarily by reverse-slip ductile shearing, ductile thrusting, and imbrication (Williams, 1991). In the Needle Mountains, southern Colorado, the structural evolution of folded quartzite and schist of the Uncompaghre Formation involved early thin-skinned, N-directed thrust faulting followed by upright folding into a large (>10 km), complex cusp- shaped fold (Harris et al., 1987; Harris, 1990). Folding was accompanied and followed by development of conjugate ductile shear zones that accommodated additional NW – SE contraction (Harris, 1990). Localized quartzite exposures across central and southern Colorado are commonly exposed in tight, upright synclinal “keels” that formed during regional subhorizontal, NW – SE crustal shortening (e.g., Gooseberry Gulch syncline; Reuss, 1974), and they are interpreted to represent the preserved roots of much larger folds. The remarkably consistent structural style, orientation, and magnitude of crustal shortening exhibited by these quartzite sequences suggest that they were deformed during the same episode of regional deformation. Although the absolute age of quartzite deformation is not well constrained, the structural styles described above are consistent with well-constrained, Mazatzal-aged deformation elsewhere across the southwestern U.S. Structures that formed during the Mazatzal Orogeny in Arizona are characterized by foreland thrust belt geometries, and estimates of shortening related to northwest-directed thrusting between 1692 – 1630 Ma range from 35% to greater than 50% (>18 km; Puls, 1986; Doe and Karlstrom, 1991; Karlstrom and Bowring, 1993). In New Mexico, N-directed crustal shortening occurred across a series of NE-striking structures after

100 deposition of 1664±3 Ma supracrustal rocks and prior to 1654±1 Ma post- kinematic plutonism (Bauer and Williams, 1994). Mazatzal-aged deformation in southern Colorado involved subhorizontal, NW – SE contraction between ca. 1658 – 1637 Ma (Shaw et al., 2001; Chapter 1 of this study).

Post-Mazatzal Orogeny (ca. 1.65 Ga) deposition

Despite the overwhelming consistency of structural styles among deformed quartzite sequences and similarities with documented Mazatzal-aged deformation across the region, there are no absolute age constraints that preclude deposition and deformation of quartzite sequences in the interval between the Mazatzal Orogeny and Mesoproterozoic (ca. 1.4 Ga) granitic magmatism. This time period coincides with a well-documented gap in U-Pb crystallization ages across the southwestern U.S. between ca. 1.60 – 1.47 Ga (Reed et al., 1993; Karlstrom et al., 2004) interpreted to represent tectonic quiescence throughout the region. Although a relatively stable tectonic environment might be favorable for the accumulation of thick sequences of passive-margin sediments, deformation of quartzites would have to have occurred prior to or during the early stages of ca.

1.4 Ga granitic magmatism or during an earlier, unrecognized deformation event. There is evidence for NW – SE shortening across the region accompanying widespread granitic magmatism at ca. 1.4 Ga, but the temporal relationship between magmatism and regional deformation does not favor the deformation of quartzites during this time. The earliest phases of Mesoproterozoic granitic magmatism began with the emplacement of a suite of

101 localized intrusions across southern Colorado between ca. 1480 – 1440 Ma (Bickford et al., 1989a). However, granites of this age are commonly undeformed and are highly discordant with respect to host-rock fabrics (Rogers et al., 2004; Bickford et al., 1989a). These early, localized pulses of Mesoproterozoic magmatism were followed by a more voluminous, widespread pulse granitic magmatism between ca. 1440 – 1420 Ma that was accompanied by amphibolite- facies metamorphism and subhorizontal NW – SE shortening and oblique shearing (Williams and Karlstrom, 1996; Graubard and Mattinson, 1990). The most substantial deformation locally occurred between ca. 1430 – 1360 Ma (see Chapter 3 of this study). Based on relationships described above, almost all of the deformation that affected quartzite exposed throughout southern Colorado occurred prior to the emplacement of ca. 1.44 – 1.43 Ga granites, essentially predating the peak of deformation. The observation that folded quartzite is locally intruded by coarse-grained granite at 1.44 Ga additionally requires that it had already been buried to mid-crustal depths prior to granite emplacement, and the structural style of deformation at ca. 1.4 Ga is inconsistent with megascopic folding that characterizes quartzite exposures across the region. Although deformation at ca. 1.4 Ga was broadly coaxial with Mazatzal-aged tectonism, much of the deformation was localized along pre-existing, crustal-scale shear zones and within the thermal aureoles of coeval granitic plutons (e.g., Nyman et al., 1994; Aleinikoff et al., 1993).

102 CONCLUSIONS

At present, there is little to no evidence throughout the southwestern U.S. for large-displacement, fold-and-thrust style deformation during the Mesoproterozoic. Furthermore, cross-cutting relationships and metamorphic assemblages described above indicate that quartzites were buried to mid-crustal depths prior to the emplacement of coarse-grained granitoids at ca. 1.44 Ga, requiring a regional tectonic event involving quartzite deformation and burial during a time interval (ca. 1.60 – 1.44 Ga) that is otherwise characterized by tectonic quiescence. Quartzite deposition in a tectonically stable environment following the Mazatzal Orogeny would be consistent with the passive margin character of the metasedimentary sequences (Soegaard and Eriksson, 1985). The extreme compositional maturity that characterizes quartzites could have been developed in this scenario through reworking of older (i.e., >1.7 Ga) detrital material. However, if quartzites did represent reworked sedimentary deposits, one might expect to find younger (i.e., Mazatzal-aged, or <1.65 Ga) detritus. As discussed above, Mazatzal-aged detritus was not recognized among the six samples analyzed during this study. This observation, coupled with regional structural arguments and deformation styles in the quartzite that are described above, leads to the conclusion that quartzite was deposited prior the Mazatzal Orogeny within a relatively narrow time window between ca. 1.70 – 1.65 Ga.

103 IMPLICATIONS

Deposition of locally-thick (1 – 2 km), supermature quartzites during a ca. 50 m.y. interval between the Yavapai and Mazatzal orogenies has important implications for the regional tectonic setting and for environmental conditions that accompanied sedimentation. Published ages from quartzite sequences in Arizona and New Mexico suggest that the onset of sedimentation began at ca. 1.7 Ga, closely following the peak of Yavapai orogenesis (e.g., Cox et al., 2002b). The quartzite sequence exposed along Blue Ridge was deposited on basement granitoids with ages of 1706 – 1698 Ma, suggesting that relatively rapid exhumation of these mid-crustal (10 – 15 km depth), Yavapai-aged rocks must have occurred prior to sedimentation. Furthermore, a widespread suite of granites emplaced between ca. 1700 – 1666 Ma across southern Colorado suggests that magmatism was ongoing at deeper crustal levels during deposition (Anderson and Cullers, 1999; Bickford et al., 1989a). These observations suggest that regional quartzite sequences were deposited in a tectonically active environment and, thus, represent syn-orogenic deposits. Deposition was closely preceded by >10 km of basement exhumation and voluminous rhyolitic volcanism and was accompanied by perhaps continued exhumation related to basin subsidence and nearly continuous granitic magmatism at deeper crustal levels. If quartzites were deposited between 1.70 – 1.65 Ga, their relatively juvenile detrital character suggests that they represent first-cycle sediments. Quartzite sequences across southern Colorado and northern New Mexico are characterized by a single, dominant detrital zircon populations that have a

104 relatively narrow range of ages (1.8 – 1.7 Ga), and the youngest grains identified only slightly predate the age of the sediments themselves. The youngest detrital ages are essentially contemporaneous with the onset of quartzite deposition. The extreme compositional purity and homogeneity of the quartzites is quite anomalous for first-cycle sediments deposited in a tectonically active environment and, thus, appears to require special environmental influences. These influences could include extreme lateritic weathering (Medaris et al., 2003); extensive eolian abrasion (Dott, 2003); diagenetic reactions that remove lithics and feldspars, increasing quartz relative to other components (Cox et al., 2002a); and/or a lack of land plants to physically bind sediment (Cox et al., 2002b). First cycle quartz arenites do presently occur in the Orinoco River basin of South America, and their compositional maturity is inferred to be related to weathering in a hot, humid climate (Johnsson et al., 1988, 1991). However, modern analogs are generally lacking, thus suggesting that environmental influences affecting quartzite deposition throughout the southwestern U.S. might have been unique to the Proterozoic.

LAURENTIAN CORRELATIVE SEQUENCES

Age constraints and depositional/structural relationships described above for Proterozoic quartzites exposed across the southwestern U.S. are remarkably consistent with other Laurentian quartzite sequences. Baraboo interval quartzites comprise seven geographically separate but lithologically and stratigraphically similar sequences exposed in the Lake Superior region of North America (Fig.

105 2.1; Dott, 1983; Dott and Dalziel, 1972). The red, supermature quartz arenites were deposited unconformably on polydeformed, Penokean-aged (1870 – 1820 Ma; Van Schmus et al., 1993) basement in a tectonically stable passive margin setting (Dott, 1983). Sedimentary characteristics of the quartzites are most consistent with deposition in a braided fluvial system (Henry, 1975; Dott, 1983), but more recent recognition of paleosols and evidence for extreme chemical maturity of the sediments suggests that unusually intense chemical weathering accompanied deposition (Medaris et al., 2003). The Baraboo interval was initially based on existing timing constraints on quartzite deposition (ca. 1750 – 1450 Ma; Dott, 1983 and references therein). More recent detrital zircon studies have helped to narrow constraints on the maximum age of sedimentation. The Baraboo Quartzite in southern Wisconsin lies nonconformably on granitic basement with ages of ca. 1750 Ma (Van Schmus et al., 2001; Van Wyck, 1995), and detrital zircon from the basal part of the quartzite yielded ages ranging from 1866 – 1691 Ma (Medaris et al., 2003). The youngest detrital zircon analyzed (1691±2 Ma; Medaris et al., 2003) is interpreted to represent the maximum age of Baraboo deposition and is consistent with reliable minimum detrital zircon ages from quartzites across the southwestern U. S. (Table 2.2, Figs. 2.8 & 2.9). Detrital zircon from correlative quartzites across the Lake Superior region have a similar range of ages (ca. 1954 – 1714 Ma; Holm et al., 1998; Van Wyck, 1995), but grains younger than ca. 1714 Ma have not been identified.

106 Absolute minimum age constraints on Baraboo interval quartzite deposition are restricted to localized cross-cutting igneous bodies emplaced during Mesoproterozoic regional magmatism. However, some of the Early Proterozoic quartzites across the Lake Superior region experienced substantial post-depositional deformation, and the northern limit of quartzite deformation corresponds well with a sharp break in cooling ages from underlying basement assemblages (Romano et al., 2000; Holm et al., 1998; Van Schmus et al., 1975). South of this thermal front, basement cooling ages were systematically reset at ca. 1630 Ma (Holm et al., 1998), and this is generally consistent with the timing of Mazatzal orogenesis documented across the southwestern U. S. (Karlstrom and Bowring, 1988, 1993). The close spatial coincidence of the deformation front in quartzites and the ca. 1630 Ma thermal front in basement assemblages is interpreted to represent thin-skinned fold-and-thrust deformation accompanied by low-grade regional metamorphism during Mazatzal-related tectonism across the Lake Superior region (Dott, 1983; Holm et al., 1998). If correct, this interpretation narrows timing constraints for quartzite deposition to between ca. 1691 – 1630 Ma, consistent with absolute and inferred constraints described above for correlative quartzite sequences across the southwestern U.S.

SUMMARY

New geochronologic studies from southern Colorado and northern New Mexico constrain the deposition of locally thick (1 – 2 km), supermature Proterozoic quartzite sequences and provide new detrital zircon age information

107 regarding their sedimentary provenance. Quartzite exposed along Blue Ridge, Colorado was deposited on granitoid basement with ages of ca. 1706 – 1698 Ma, and cross-cutting pegmatite dikes require that quartzite deposition and two episodes of deformation occurred prior to ca. 1436 Ma. Detrital zircon geochronology from five quartzite localities reveals that the metasedimentary sequences are characterized by Paleoproterozoic-aged detritus that define a relatively narrow age range (1.80 – 1.70 Ga). Although Archean-aged grains were present in most cases, they represented a volumetrically minor component of the detrital population. Peak detrital ages among all quartzites sampled similarly defined a narrow range of ages (1.76 – 1.70 Ga) that generally reflect the average age of surrounding basement exposures. The youngest detrital zircon grains (ca. 1.70 Ga U-Pb ages) confirm the maximum age of sedimentation suggested by underlying basement granitoids at Blue Ridge. No Mazatzal-aged (ca. 1.65 Ga) detrital zircon was recognized among all samples analyzed during this study. These new results require that quartzite deposition occurred after the ca. 1.70 Ga Yavapai Orogeny but prior to regional granitic magmatism at ca. 1.43 Ga. These new constraints define a relatively broad time window (ca. 250 m.y.) for deposition. However, regional structural arguments, the deformational style of quartzite, and a lack of Mazatzal-aged detritus suggest that quartzites were deposited during a much narrower (ca. 50 m.y.) time window between the Yavapai and Mazatzal orogenies. This interpretation requires that quartzite deposition occurred in a tectonically active environment involving rapid exhumation of mid-crustal rocks and contemporaneous granitic magmatism at

108 deeper crustal levels. Detrital zircon populations are characterized by a relatively narrow age range (1.80 – 1.70 Ga) that only slightly predates the age of quartzite deposition, suggesting that quartzites represent first-cycle sediments. Resolving the syn-orogenic tectonic setting with the extreme compositional maturity characterizing the first-cycle quartzites would require enhanced chemical weathering during deposition that was likely related to anomalous environmental influences (Medaris et al., 2003; Dott, 1983). Quartzite sequences with similar characteristics and age constraints occur throughout the southwestern U.S. and in the Lake Superior region of North America, suggesting that the metasedimentary sequences might represent an important episode of broadly contemporaneous sedimentation along the entire Laurentian margin.

109 Chapter 3. Contrasting structural styles at ca. 1.4 Ga across the Wet Mountains, Colorado: Implications for models for intracontinental tectonism in the southern Rocky Mountains

ABSTRACT

New field and geochronologic studies in the Wet Mountains, Colorado, reveal that contrasting structural styles from N – S across the range were coeval at ca. 1.4 Ga. Deformation in the northern part of the range produced or tightened NE-striking, subvertical fabrics and folds between ca. 1430 – 1404 Ma. A new U-Pb zircon age from a deformed pegmatite dike within the N-striking Five Points Gulch shear zone indicates that at least one episode of deformation occurred after 1430+5/-3 Ma, accommodating reverse-sense (E-side-up) displacement. Deformation in the central and southern part of the range produced moderately- to shallowly-dipping, penetrative gneissic fabrics and accompanied two episodes of voluminous granitic magmatism at 1435±4 Ma and 1390±10 Ma. The N – S transition from subvertical fabrics and localized deformation to penetrative, moderately-dipping fabrics is mirrored by a change in the style of plutonism from discrete magmatic bodies in the north to a distributed, concordant magmatic “framework” in the south. Deformation at ca. 1.4 Ga throughout the range is broadly kinematically compatible with subhorizontal, NNW – SSE shortening, but contrasting, yet coeval, styles of deformation suggest a profound structural discontinuity in the middle crust, perhaps corresponding with the brittle-ductile transition that was localized by emplacement of voluminous, A-

110 type granitic magmas (Shaw et al., 2005). The central and southern Wet Mountains are interpreted to represent relatively deep levels of crustal exposure, and moderately-dipping fabrics are interpreted to reflect subhorizontal ca. 1.4 Ga crustal flow. Weak, flowing lower crust beneath relatively strong, brittle upper crust is consistent with models for the development of intracontinental orogenic plateaus. Futhermore, moderately-dipping fabrics, together with the voluminous, concordant framework of contemporaneous magmas, might represent a deeply- exhumed analog to mid-crustal melt-rich layers inferred in modern intracontinental orogenic settings.

INTRODUCTION

The regional context and tectonic setting for widespread Mesoproterozoic (ca. 1.4 Ga) granitic magmatism across the southwestern United States has long been a subject of debate. Early tectonic models centered on the perceived “anorogenic” character of the intrusions and suggested continental rifting or extension (e.g., Anderson, 1983). More recent recognition of associated amphibolite-facies metamorphism and localized yet kinematically consistent regional deformation, including reactivation of lithosphere-scale shear zones (e.g., Nyman et al., 1994; Shaw et al., 2001), led to the development of models involving thermally (Ferguson et al., 2004) or tectonically driven processes (Nyman et al., 1994). One key approach to testing these various models involves carefully constraining both the spatial and temporal relationship between magmatism and deformation and the kinematics and style of deformation. This

111 approach not only provides the most precise constraints on a local tectonic history but also permits more confident correlation and evaluation of fabric elements across a broad region in a coherent tectonic framework. Rocks in the Wet Mountains, Colorado, experienced a ca. 1.4 Ga thermal culmination (Shaw et al., 2005) related to voluminous magmatism (>1000 km3) coeval with associated high-temperature metamorphism and deformation. Rocks across the range are some of the deepest levels of crustal exposure across the southwestern U.S., and structural and metamorphic observations suggest that the range also exposes an oblique cross-section through Proterozoic crust profoundly affected by ca. 1.4 Ga tectonothermal events (Siddoway et al., 2000). From north to south across the range, this oblique crustal section is characterized by a change in structural style from upright, open folding and subvertical, localized deformation to subhorizontal, penetrative flow fabrics. Furthermore, these contrasting deformation styles are mirrored and perhaps directly influenced by a change from north to south in the style of magmatism from discrete intrusive bodies to a voluminous, concordant framework of dikes and sills that profoundly altered the rheology of host-rock gneisses. Taken together, these contrasts are interpreted to result from a difference in unroofing from 5 – 10 km in the north to >20 km in the south. Strongly localized, subvertical deformation with discrete, localized magmatism is interpreted to represent shallower crustal levels, whereas penetrative crustal flow amid a framework of voluminous granite is interpreted to represent deeper levels (20 – 25+ km) of crustal exposure. The inferred transition between the two domains is interpreted to represent a fundamental rheological

112 contrast in the crust approximating the brittle-ductile transition developed at ca. 1.4 Ga. This transition was likely localized in the crust at a level of neutral buoyancy where granitic magmas pooled and were locally tapped by through- going crustal weaknesses (e.g., shear zones) and emplaced at shallower levels (Shaw et al., 2005). New field and geochronologic studies were undertaken in the Wet Mountains, Colorado, to constrain the history of this mid-crustal transition. Samples were collected along a N – S transect across the range to constrain the temporal kinematics of deformation both above and below the inferred transition and the temporal and spatial relationship between deformation and magmatism. New data reveal evidence for long-lived, penetrative crustal flow in the southern Wet Mountains and suggest that the mid-crustal structural discontinuity developed by ca. 1430 Ma and persisted 70 m.y. until ca. 1360 Ma. These new results not only have direct implications for tectonic models for ca. 1.4 Ga magmatism and deformation across the southwestern U.S. but also provide new insights into fundamental tectonic processes including the mechanics of intracratonic deformation and the development of orogenic plateaus.

GEOLOGIC SETTING

Precambrian mid-crustal exposures across the southwestern U.S. comprise metavolcanic rocks, metasedimentary rocks, and mafic and granitoid plutons that were accreted to the southern margin of Laurentia between 1.8 – 1.6 Ga (Condie, 1982; Karlstrom and Bowring, 1988). The Yavapai Province represents the

113 earliest phase of Paleoproterozoic accretion during which a complex collage of juvenile arc terranes were added between 1.78 – 1.70 Ga along a belt stretching from Colorado to Arizona. This phase of southward arc growth and collision culminated with the ca. 1.7 Ga Yavapai Orogeny (Karlstrom and Bowring, 1988; 1993). Following stabilization of the newly-formed Yavapai crust, which involved voluminous post-orogenic granitoid magmatism and erosional unroofing to 10 – 15 km depth followed by regional quartzite deposition (Anderson and Cullers, 1999; Williams et al., 2003), the Mazatzal Province was accreted to the southern margin across New Mexico and Arizona during the 1.65 – 1.63 Ga Mazatzal Orogeny (Silver, 1965; Karlstrom and Bowring, 1988). Mazatzal-aged foreland deformation also propagated northward into older, Yavapai-aged crust of Colorado, and the “Mazatzal Front” represents the approximate northern extent of these effects (Fig. 3.1; Karlstrom and Bowring, 1993; Karlstrom and Daniel, 1993). A discrete regional structure forming the boundary between these two provinces has not been identified, and instead it appears to be a broad, complex, low-angle deformation zone characterized by refolded thrusts and duplexes (Shaw and Karlstrom, 1999). Following the Mazatzal Orogeny, the region experienced a 200 m.y. magmatic and tectonic lull followed by a widespread regional episode of granitic magmatism at ca. 1.4 Ga. Magmatism was accompanied by coeval mafic diking, high-temperature/low-pressure metamorphism and deformation. Rocks of this age presently account for nearly 20% of all Precambrian exposures across the region

114

Figure 3.1. Regional (southwestern U.S.) index map. Precambrian exposures (grey) and ca. 1.4 Ga granites (red) are emphasized. Proterozoic crustal provinces, inferred boundaries and/or transition zones, and approximate age ranges are shown (Condie, 1986; Bennett and DePaolo, 1987; Karlstrom and Bowring, 1988; Wooden et al., 1988; Wooden and DeWitt, 1991). Regional qualitative strain ellipse inferred from structural data from the southern Rocky Mountains and Arizona (Graubard and Mattison, 1990; Kirby et al., 1995; Nyman and Karlstrom, 1997; Shaw et al., 2001; Selverstone et al., 2000).

115 (Fig. 3.1). This voluminous pulse of magmatism fundamentally reworked much of the juvenile continental lithosphere across southern Laurentia (Anderson, 1987). Granites of this age have distinct A-type geochemical characteristics (Loiselle and Wones, 1979; Anderson, 1983) that are typically associated with extensional environments. However, widespread evidence for contemporaneous deformation suggests that they were emplaced during a regional tectonic event, perhaps related to renewed convergence along a distal southern margin (Nyman et al., 1994). Although isotopic and geochemical data suggest that crustal growth and accretion continued along the southern margin of Laurentia during the Mesoproterozoic (1.5 – 1.3 Ga; Bennett and DePaolo, 1987; Patchett and Ruiz, 1989; Barnes et al., 1999), no surface exposures of direct evidence of a margin have been identified.

PROTEROZOIC LITHOLOGIES AND STRUCTURAL ELEMENTS OF THE WET MOUNTAINS

The Wet Mountains comprise a large (100 km x 30 km) block of nearly continuous Precambrian exposure segmented locally by Phanerozoic brittle faults

(Fig. 3.2). The range lies within the Yavapai province south of the Mazatzal deformation front (Fig. 3.1; Shaw and Karlstrom, 1999). Basement lithologies in the northern Wet Mountains are characterized by a diverse assemblage of supracrustal metavolcanic and metasedimentary rocks dominated by quartzose and quartzo-feldspathic gneisses but also including abundant schist, calc-silicate gneiss, mafic gneiss, and amphibolite. Map-scale lithologic units range in thickness from tens to hundreds of meters, and contacts are commonly sharp but 116

Figure 3.2. Generalized Precambrian geology of the Wet Mountains. Areas sampled for new geochronology and summary of new U-Pb zircon ages are indicated along with a summary of published U-Pb ages.

117 are locally gradational across distances of a few meters. The best exposures are accessed along an E – W canyon cut across the northern part of the range by the Arkansas River (Arkansas River Gorge = ARG; Fig. 3.2). Basement exposures throughout the central and southern Wet Mountains are less compositionally diverse and comprise interlayered quartzose and quartzo-feldspathic gneiss, and amphibolite. Local exposures of schist, marble, and calc-silicate gneiss are rare. In general, metamorphic grade increases from north to south across the range. Whereas ARG exposures experienced peak metamorphic conditions of greenschist- to amphibolite-facies (Siddoway et al., 2000), rocks in the central and southern Wet Mountains most typically exhibit upper-amphibolite- to granulite- grade textures and mineral assemblages (Boyer, 1962; Lanzirotti, 1988), and migmatitic rocks are common (Boyer, 1962). These gneissic rocks are cut by numerous granitoid intrusions that comprise two general age groups of plutons, as first delineated through geochronologic studies of Bickford et al. (1989a). The older group includes late- to post-Yavapai granitoids exposed along the ARG that range in age from ca. 1705 – 1663 Ma (Fig. 3.2; Bickford et al., 1989a). The younger group comprises a variably deformed suite of Mesoproterozoic granites emplaced between 1474 – 1361 Ma (Fig. 3.2; Bickford et al., 1989a). Siddoway et al. (2000) recognized three phases of magmatism across the central and southern Wet Mountains based on field observations of textural and cross-cutting relationships. The oldest granitoids (G1) are commonly coarse-grained to K-feldspar-megacrystic bodies that are variably discordant with respect to host-rock gneiss foliation and display a

118 range of deformational characteristics. G1 granitoids are generally Paleoproterozoic in age and include foliated tonalite and granodiorite of the ca. 1705 Ma Twin Mountain and Crampton Mountain plutons and undeformed granodiorite of the ca. 1663 Ma Garell Peak pluton (Fig. 3.2; Bickford et al., 1989a). Second-phase granitoids (G2) are similarly coarse-grained to K-feldspar megacrystic, but are more commonly concordant with respect to host-rock gneiss foliation and are locally strongly deformed. These granitoids are compositionally and texturally correlated with Mesoproterozoic (ca. 1.4 Ga) granitic intrusions across the region and include the foliated quartz monzonite of the ca. 1439 Ma Oak Creek pluton (Fig. 3.2; Bickford et al., 1989a) and a newly-dated suite of ca. 1434 Ma gneissic granite sills exposed across the southern part of the range (see below). The youngest suite of granitoids (G3) is characterized by fine-grained granitic sills that are locally discordant with respect to host-rock gneiss foliation but commonly contain a well-developed, metamorphic foliation that is parallel with the surrounding host-rock and granite fabric. New geochronology described below indicates that they were emplaced at ca. 1390 Ma. In general, the style of ca. 1.4 Ga magmatism changes dramatically from north to south across the range. To the north, granitoids are commonly exposed as discrete, map-scale plutons or stocks with sharp contacts (e.g., West McCoy Gulch pluton; Fig. 3.2). Although similar map-scale intrusions also occur in the central and southern Wet Mountains (e.g., Oak Creek pluton, San Isabel pluton; Fig. 3.2), basement exposures across this part of the range are locally dominated by a concordant network of G2 and G3 granitoid sills and dikes. These igneous

119 bodies were emplaced as a distributed magmatic framework that only locally makes up 100% of outcrops up to a few hundred metes in size. This relationship makes it difficult to map the occurrence and extent of ca. 1.4 Ga intrusive rocks across much of the range and helps to explain the general paucity of map-scale intrusive bodies represented in Figure 3.2. The transition between discrete plutonism and diffuse “framework” magmatism is difficult to define but might occur across a kilometer-wide zone of voluminous, presumably ca. 1.4 Ga, granite and pegmatite south of the ARG (Siddoway et al., 2000). The change in the style of magmatism is mirrored by a change in structural style and fabric orientation from north to south across the range. In the northern Wet Mountains, Proterozoic exposures are generally characterized by steeply-dipping to subvertical fabrics, and deformation is localized in discrete structural domains. Although deformation is broadly compatible with NNW- to NW-directed shortening, two domains across the ARG, the Texas Creek and Parkdale domains, were delineated based on characteristic fabric orientations, deformation styles, and/or fabric intensity. These domains are juxtaposed across the 2 – 5 km thick Five Points Gulch shear zone, a NNW-striking, subvertical zone that locally contains mylonitic fabrics. Subvertical fabrics and localized deformation across the ARG contrast sharply with moderately-dipping fabrics and penetrative deformation that characterize exposures throughout the central and southern Wet Mountains. These various structural elements and styles of deformation are described from N – S across the range in the following sections.

120 Northern Wet Mountains (ARG) structural elements

Fabric and structural elements in the Texas Creek domain (Plate 2) were described by Siddoway et al. (2000) and are briefly summarized herein. Two phases of deformation affected a compositionally diverse assemblage of gneisses.

Early deformation (D1) produced a penetrative crenulation cleavage (S1) that disrupts compositional layering and inferred bedding in the dominantly metasedimentary package. D1 was accompanied by metamorphism (M1) during which large (10 – 20 cm) cordierite porphyroblasts grew throughout pelitic schists. Siddoway et al. (2000) interpreted D1 and M1 to have occurred at ca. 1.66 Ga, broadly synchronous with emplacement of the Garell Peak pluton (1663±4 Ma; Bickford et al., 1989a) <1 km to the west (Plate 2). A second deformation event (D2) involved folding of S1 into a subparallel series of upright, kilometer- scale, E-trending folds (F2) with moderately-plunging axes (average orientation = 45°/081; Siddoway et al., 2000). These folds are interpreted to have formed during subhorizontal, NNW-directed shortening (Siddoway et al., 2000). The deflection of fabrics in folded gneisses around the 1474±7 Ma West McCoy Gulch pluton (Bickford et al., 1989a) and presence of ca. 1.43 – 1.42 Ga monazite inclusions in cordierite porphyroblasts in F2 fold limbs (Siddoway et al., 2002) suggest that D2 was broadly coeval with G2 granitic magmatism. Megascopic, W-trending folds in the Texas Creek domain are sharply truncated to the east by the subvertical, NNW-striking Five Points Gulch shear zone (FPSZ, Plate 2; Siddoway et al., 2000). The FPSZ is a 2 – 5 kilometer wide

121 zone exhibiting homogeneous, high-temperature fabrics that affected a compositionally and texturally uniform suite of garnet-K-feldspar-biotite-quartz- plagioclase gneisses. These gneisses are locally cut by pods or lenses of amphibolite up to tens of meters long that sharply truncate the gneissic foliation but also contain a well-developed foliation that is parallel with the shear zone fabric. The dominant fabric across the shear zone (Ssz) strikes NNW and dips steeply ENE (average orientation = 335, 63° E; Plate 2A). A well-developed mineral lineation defined by porphyroblastic sillimanite (Givot and Siddoway, 1998) generally plunges moderately NNE (average orientation = 45°/015; Plate 2B) but is locally steep to subvertical. The metamorphic assemblage present across the FPSZ indicates peak metamorphic conditions of >700°C and 500 MPa accompanying deformation (Givot, 1998). Kinematic indicators across the shear zone include ductile shear bands, en- echelon tension gash arrays, and, locally, asymmetric tails on garnet. Many of these indicators show a close association with pegmatitic melts, suggesting a genetic relationship between magmatism and at least one episode of movement along the shear zone. The dominant sense of displacement is east side up, consistent with an increase in metamorphic grade from west to east across the FPSZ (Siddoway et al., 2000). Kinematic indicators and asymmetric folds in parts of the shear zone with steeply-plunging lineations agree with reverse-sense displacement (Fig. 3.3A & B), but more localized asymmetric fabrics indicate a significant component of sinistral-oblique movement during deformation. Complex local kinematic relationships are likely related to multiple phases of

122 movement along the shear zone, but cross-cutting relationships between asymmetric fabrics are rare. Because the FPSZ truncates F2 folds of the adjacent Texas Creek domain, the major phase of shear zone development is a third phase of deformation (D3) that occurred during the Mesoproterozoic (Siddoway et al., 2000). Siddoway et al. (2000) hypothesized that the N – S orientation of the shear zone, atypical of the dominant NE-striking tectonic grain across much of the region, might represent part of a conjugate shear system developed during NW- directed shortening. Alternatively, the shear zone might reflect the concentration of strain along a pronounced rheological anisotropy where layered gneisses lie adjacent to the large (>100 km2), relatively massive ca. 1.7 Ga Crampton Mountain batholith (Plate 2). The Parkdale domain comprises the easternmost 10 kilometers of the ARG and is bounded on the east by high-angle Phanerozoic brittle faults related to the Tertiary Parkdale graben (Plate 2). Exposures throughout the domain are dominated by tonalite, quartz diorite, and granodiorite of the Crampton Mountain pluton (1705±8 Ma; Bickford et al., 1989a) and a thin (<1 km) panel of host-rock gneisses compositionally similar to those in the adjacent FPSZ. The boundary between the shear zone and Parkdale domain is gradational over a distance of a few hundred meters and represents the transition from the NNW-striking shear zone fabric (Ssz) to subvertical, NE-striking fabrics (Plate 2). This fabric transition spatially coincides with the western margin of the 150 km2 pluton, and the orientation of foliation in host-rock gneisses generally reflects the map-scale geometry of the intrusion margin (Plate 2). These relationships further suggest

123

Figure 3.3. Field photographs from the eastern Arkansas River Gorge. A) Cross- sectional view looking to south of deformed pegmatite dike (sample J01-FP1, Plate 2 for location) from within the Five Points Gulch shear zone (FPSZ). Overall shear sense across the zone is E-side-up. B) Close-up of shear bands that offset the pegmatite dike shown in (A). C) View down on deformed quartz diorite of the Crampton Mountain batholith (sample J03-TM1, Plate 2 for location) with shear band and elongate, mafic xenolith.

124 that the geometry and location of the intrusive body might have fundamentally controlled the orientation of fabrics across the Parkdale domain and FPSZ. Plutonic rocks across the Parkdale domain contain a NE-striking, subvertical solid-state foliation (average orientation = 235, 87° NW; Plate 2C) defined by biotite and amphibole and locally enhanced by dynamically- recrystallized feldspar (plagioclase and K-feldspar) and quartz. Abundant mafic xenoliths are also flattened parallel with the surrounding fabric (aspect ratios up to 30:1). The average foliation orientation in gneissic host rocks varies throughout small, localized exposures in the western part of the domain. A well-developed biotite and/or amphibole mineral lineation occurs throughout igneous exposures and plunges moderately to steeply NE (average orientation = 62°/045; Plate 2D). Asymmetric fabrics including composite (C-S) foliation and ductile shear bands indicate dominantly reverse-sense (NW-side-up) displacement across the southern margin of the Crampton Mountain pluton with a component of sinistral offset. The timing of Parkdale domain deformation relative to FPSZ development is difficult to constrain due to the transitional nature of the domain/shear zone boundary, but compatible kinematics suggest that deformation could have been coeval during the Mesoproterozoic.

Central and southern Wet Mountains structural elements

In contrast to the ARG exposures, fabrics across the central and southern Wet Mountains are remarkably consistent over a large area (>100 km2) and are prevalent throughout both host-rock gneisses and multiple intrusive granitic

125 phases (Fig. 3.4). Average foliation in the central Wet Mountains strikes E – W with moderate to shallow NNW dips (average orientation = 270, 50°N; Fig. 3.5A) and is commonly accompanied by a well-developed, N-plunging mineral lineation defined by biotite (average orientation = 52°/003; Fig. 3.5B). The foliation is both folded by and axial planar to macroscopic (20 cm – 1 m) to megascopic (>1 km), tight-to-isoclinal folds of gneissic layering and granitic sills. This relationship suggests that foliations across the central and southern Wet Mountains represent a composite fabric formed during multiple episodes of broadly coaxial deformation. Lanzirotti (1988) recognized three phases of deformation (D1 – D3) involving isoclinal folding and foliation development in the central Wet Mountains. In the southern Wet Mountains, there is some discrepancy between foliation orientations in G2 and G3 granitic intrusions and gneissic host rocks, respectively. Gneisses have an average foliation that strikes E – W and dips moderately N (average orientation = 283, 56°N; Fig. 3.6A), consistent with the dominant fabric across much of the range. Although the mineral lineation contained within both gneissic and granitic rocks is essentially parallel (Fig. 3.6B & D), G2 and G3 granitic sills contain a well-developed, locally gneissic foliation that strikes ENE – WSW and dips moderately NNW (average orientation = 243, 43°N; Fig. 3.6C). The gneissic host-rock foliation is interpreted as a long-lived, composite fabric perhaps developed during the Paleoproterozoic. Although new U-Pb zircon data suggests that metamorphic recrystallization of this fabric occurred during the Mesoproterozoic (J01-WC1, see below), the pre- existing

126

Figure 3.4. Field photographs of widespread, penetrative, moderately-dipping fabrics in the central (A) and southern (B) Wet Mountains. View is to the west (A) and northeast (B), respectively. The prominent pinnacles in the center of photograph (A) are fine-grained, G3 granitic sills. “The Wall” in photograph (B) is the large, vertical face from which geochronology samples from the southern Wet Mountains were collected (indicated by red stars).

127

Figure 3.5. Generalized geologic map of the North Hardscrabble Creek area, central Wet Mountains. The Rattlesnake Gulch outcrop that was sampled for geochronology is indicated by the arrow. Foliation (plotted as poles) and lineation data are plotted on lower-hemisphere, equal- area stereonet diagrams, and rock types from which measurements were taken are indicated to the left of the diagrams. Structural data along Colorado Highways 96 and 165 is from this study; structural data north of the highway is from Siddoway et al. (2000). Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003). 128

Figure 3.6. Foliation map for Precambrian exposures in the Bear Creek/Williams Creek area of the southern Wet Mountains. “The Wall” outcrop from which geochronology samples were collected is indicated (red star). Geologic contacts are not shown because of intricately intermingled character of granites and basement gneisses in individual outcrops. Foliation (plotted as poles) and lineation data are plotted on lower-hemisphere, equal-area stereonet diagrams. Rock types from which measurements were taken are indicated to the left. All structural data is from this study, and average orientations were calculated using GEOrient 9.1 (Holcombe, 2003). 129 anisotropy was likely too pronounced to be reoriented during renewed crustal shortening. Warmer, relatively isotropic intrusive rocks are interpreted to have acted as strain guides during Mesoproterozoic deformation, and, thus, their foliation is expected to more accurately reflect the orientation of shortening and suggests NNW- directed shortening during this time.

U-PB ZIRCON AND TITANITE GEOCHRONOLOGY

A suite of samples was collected for U-Pb geochronology along a N – S transect across the Wet Mountains to constrain the timing of magmatism, deformation, and metamorphism. Samples were carefully chosen to precisely constrain the development of contrasting structural styles described above. Isotopic data are presented in Table 3.1, and associated concordia diagrams are contained in Figures 3.7, 3.9, and 3.10. Zircon and titanite fractions were hand picked, examined using a petrographic microscope, in some cases characterized by cathodoluminesence, extensively abraded, and then subjected to a final optical re-evaluation before analysis. For complete analytical methods refer to Appendix 1. Results of this work agree well with existing geochronology across the range (Fig. 3.2, Bickford et al., 1989a) and demonstrate that contrasting styles of deformation across the Wet Mountains were essentially coeval at ca. 1430 Ma and accompanied a widespread regional pulse of A-type granitic magmatism (Reed et al., 1993). In the southern Wet Mountains, penetrative deformation was accompanied by high-temperature metamorphism at ca. 1436 Ma, and

130 deformation continued intermittently until emplacement of the San Isabel granite at ca. 1361 Ma (Bickford et al., 1989a).

Northern Wet Mountains – Arkansas River Gorge

With timing relationships of deformation and metamorphism constrained at multiple locations in the Texas Creek domain (Siddoway et al., 2002; Siddoway et al., 2000), sampling for new U-Pb geochronology was concentrated in the FPSZ and Parkdale domain to constrain the timing of shear zone development and fabric reactivation in the eastern ARG (Plate 2).

DEFORMED PEGMATITE DIKE (J01-FP1). This sample was collected from a 20 – 30 cm thick, subvertical pegmatite dike on the south side of Colorado state highway 50 0.5 km west of Sheep Basin. The dike is texturally and compositionally correlated with a suite of subvertical dikes emplaced parallel with the NNW-striking fabric characterizing the FPSZ (Ssz; Plate 2A). These intrusions both follow and sharply cut across the shear zone fabric and are interpreted to be coeval with at least one major phase of shear zone deformation

(D3). The sampled dike locally truncates the shear zone fabric but is deformed by a series of en-echelon, E-dipping shear bands. Offset of pegmatite blocks indicates reverse-oblique kinematics (Fig. 3.3A & B), consistent with E-side up displacement across the FPSZ (Siddoway et al., 2000). The sampled pegmatite dike yielded a single population of tan to brown, euhedral to subhedral zircon with prismatic faces consistent with an igneous

131 origin. Optical analysis revealed that most of the grains were characterized by dark brown, xenocrystic cores surrounded by light tan to clear, euhedral tips (Fig. 3.7A inset). Tips were mechanically separated from core material using tweezers, and the respective parts of the grains were dissolved and analyzed separately to determine the ages of zircon comprising both cores and rims. Two fractions of dark brown zircon core material (Z1 and Z5) overlap concordia with 207Pb/206Pb ages of 1451 and 1449 Ma, respectively (Table 3.1; Fig. 3.7A). Three fractions of zircon tip material (Z2, Z3, and Z6) define a line with intercepts of 1430+5/-3 Ma and 228±300 Ma (Fig. 3.7A). Zircon cores are interpreted to represent grains that were inherited from nearby Mesoproterozoic granitic plutons (e.g., Hindman Gulch pluton; Plate 2), and euhedral, clear zircon tips are interpreted to have grown during crystallization of the pegmatite dike. The lower intercept reflects more recent Pb loss likely related to Ancestral Rockies or Laramide tectonism. The crystallization age not only constrains a major phase of movement along the FPSZ, but it also corresponds with a regional suite of A-type granitic plutons emplaced between ca. 1440 – 1430 Ma (Reed et al., 1993).

FOLIATED QUARTZ DIORITE (J03-TM1). This sample of foliated quartz diorite was collected in the Parkdale domain from the southwestern part of the Crampton Mountain pluton (Plate 2) 1 km west of the dated locality of Bickford et al. (1989a). Tonalite, quartz diorite, and granodiorite comprising the pluton were emplaced at 1705±8 Ma (Bickford et al., 1989a) and are the oldest phases of a more regional suite of variably-deformed granitoid intrusions interpreted to

132 represent the waning stages and post-collisional phase of Yavapai orogenesis (Anderson and Cullers, 1999). The sample contains a subvertical, NE-striking foliation (Fig. 3.3C) defined by dynamically recrystallized feldspar and biotite and a steep, NE-plunging biotite mineral lineation. Asymmetric mineral fabrics and numerous ductile shear bands throughout the sampled outcrop suggest reverse-oblique (NW-side-up) kinematics. Although these fabrics are consistent with the NE-trending regional tectonic grain developed during the Yavapai Orogeny (Karlstrom and Humphreys, 1998), various workers have demonstrated that similarly-oriented fabrics were reactivated across the region during broadly coaxial Mazatzal and/or ca. 1.4 Ga tectonism (e.g., Shaw et al., 2001; McCoy et al., 2005). Because the emplacement age of the Crampton Mountain pluton was already well established, this sample was collected for titanite to determine whether or not metamorphism related to reactivation of the foliation occurred during either of these younger tectonic pulses. The sample yielded abundant dark brown, angular fragments of titanite. In thin section, titanite occurs in small (sub-millimeter) clusters that are elongate parallel with the dominant foliation. Three fractions (T1 – T3) analyzed have 207Pb/206Pb ages ranging from 1443 – 1404 Ma (Table 3.1), and fractions T2 and T1 overlap concordia at 1422 Ma and 1404 Ma, respectively (Fig. 3.7B). These ages are interpreted to represent metamorphic recrystallization of titanite accompanying reactivation and enhancement of the NE-striking foliation, and the range of ages are interpreted to reflect pulses of metamorphism occurring throughout an otherwise protracted tectonothermal event.

133

Figure 3.7. U-Pb concordia diagrams for samples from the Arkansas River Gorge, northern Wet Mountains. Ages are determined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

134

135

136

137 Central Wet Mountains – Rattlesnake Gulch

Sampling for U-Pb geochronology in the central Wet Mountains was concentrated along a single large roadcut on the northern side of Colorado Highway 69 3 km east of McKenzie Junction (Figs. 3.5 & 3.8). Geographic coordinates of the outcrop are included in Table 3.1. Host-rock amphibolite and felsic gneiss and a thick, fine-grained granite sill were collected to determine the protolith ages for basement gneisses, the timing of granitic magmatism, and the timing of deformation and metamorphism in this part of the range.

AMPHIBOLITE (J01-RG3A). This sample of amphibolite comprises the main compositional component of basement gneisses surrounding the fine-grained granitic sill described below (sample J01-RG1; Fig. 3.8), and it was collected 2.5 meters beneath the intrusion margin to determine the protolith age of gneisses in the central Wet Mountains and/or the timing of peak metamorphism in this part of the range. The medium- to coarse-grained amphibolite contains a well-developed, penetrative amphibole foliation that is parallel with compositional and gneissic layering across the outcrop. The fabric strikes E – W, dips moderately N, and is accompanied by a moderately NNE-plunging amphibole lineation (Fig. 3.5B). This sample yielded a relatively simple population of subhedral to anhedral, colorless zircon with some euhedral, prismatic grains. Cathodoluminescence (CL) imaging reveals diverse internal characteristics ranging from concentric growth zonation to complex patchy and sector zoning patterns (Fig. 3.9A insets), and most grains exhibited overgrowths interpreted to represent secondary, and

138 likely metamorphic, zircon growth. Six zircon fractions plot within an envelope of ages defined by an older reference line with intercepts of 1680 Ma and 550 Ma and a younger reference line with intercepts of 1436 Ma and 0 Ma (Fig. 3.9A). Two fractions (Z1 and Z3) define the older line and are interpreted to record crystallization of the mafic protolith, an age that is in agreement with the oldest zircon fractions from both the felsic gneiss and granitic sill sampled from the same outcrop. The lower intercept of this reference line is interpreted to represent more recent Pb loss, probably related to Ancestral Rockies tectonism. One fraction (Z2) overlaps concordia with a 207Pb/206Pb age of 1436 Ma (Table 3.1), and this age is interpreted to represent new metamorphic zircon growth contemporaneous with metamorphism in the southern Wet Mountains (J01-WC1, see below). The three remaining fractions (Z4 – Z6) have 207Pb/206Pb ages ranging from 1605 – 1560 Ma and plot just below a mixing line between the upper intercepts of the two reference lines (Fig. 3.9A). These zircon fractions are interpreted to reflect mixing of older (Paleoproterozoic), protolith-aged cores with varying amounts of younger (Mesoproterozoic) metamorphic zircon overgrowth. This sample also yielded abundant pale yellow, angular titanite fragments. In thin section, titanite occurs in elongate, millimeter-size clusters parallel with the dominant amphibole fabric. Three fractions (T1 – T3) overlap concordia with an average 207Pb/206Pb age of 1390±7 Ma (Fig. 3.9A; Table 3.1). Whereas earlier metamorphism of the amphibolite is recorded by ca. 1436 Ma zircon

139

Figure 3.8. Field photographs (cross-sectional view looking to north) from the Rattlesnake Gulch area, central Wet Mountains. The photo panorama shows the locations of the three lithologies sampled for U-Pb geochronology. A) Isoclinally folded dike beneath fine-grained granite sill contains an axial-planar biotite foliation. B) Granite-filled boudin neck in host-rock amphibolite (center of photo) and concordant nature of the granite sill and its apophyses.

140

Figure 3.9. U-Pb concordia diagrams for samples from the Rattlesnake Gulch area, central Wet Mountains. Ages are determined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details. 141 growth, this younger age is interpreted to reflect metamorphic recrystallization of titanite possibly accompanying local fabric reactivation.

FELSIC GNEISS (J01-RG2). Quartz-rich felsic gneiss makes up part of the host rock to the fine-grained granitic sill described above, and this sample was collected from 2 meters below the intrusion (Fig. 3.8) to determine the protolith age for gneisses in this part of the Wet Mountains. The felsic gneiss contains a well-developed, penetrative biotite and quartz foliation that is parallel with gneissic layering across the outcrop. This sample yielded zircon morphologies that range from euhedral to subhedral and equant to slightly elongate with some preserved prismatic faces. Euhedral to subhedral, prismatic morphologies are generally consistent with an igneous origin, but a number of anhedral zircon fragments were present as well. Cathodoluminescence (CL) imaging reveals internal characteristics similar to zircon from sample J01-RG1 (see below) with bright, concentrically-zoned cores variably surrounded by dark, featureless overgrowths interpreted to represent secondary zircon growth. Three fractions were analyzed and have 207Pb/206Pb ages ranging from 1691 – 1643 Ma (Table

3.1; Fig. 3.9B). The oldest of these fractions (Z1) is interpreted to represent the approximate protolith age for basement gneisses in the central Wet Mountains, and the two younger fractions (Z2 and Z3) are interpreted to represent metamorphic recrystallization of zircon. Testing these interpretations would require additional analyses, but the observation that these data do not reveal any

142 Mesoproterozoic zircon growth or recrystallization is the most important conclusion that can be drawn at this time.

FINE-GRAINED, FOLIATED GRANITE SILL (J01-RG1). This sample is from a 5 – 10 meter thick, fine-grained granite sill that was emplaced into interlayered, amphibolite-facies felsic gneiss and amphibolite (Fig. 3.8). The sill contains a well-developed biotite foliation that strikes E – W and dips moderately N and a moderately NNE-plunging mineral lineation defined by biotite and quartz (Fig. 3.5). Locally, the granite sharply cuts compositional layering and a well- developed gneissic foliation in the host gneisses, but thin (5 – 10 cm) granitic apophyses are tightly folded (Fig. 3.8A). Fold hinge lines are sub-parallel with the mineral lineation, and axial planes are parallel with the dominant gneissic fabric. Although symmetrical folds and flattening fabrics dominate across the outcrop, the granite also contains asymmetric mineral fabrics indicating reverse- sense (top-up-to-the-SSE) kinematics. Previous attempts to date this sill were unsuccessful (Siddoway et al, 2000), and two separate 30 – 40 kg samples were required to yield enough zircon for this study. The granite yielded a small but relatively simple population of euhedral to subhedral zircon with some large, blocky, anhedral fragments. Cathodoluminescence (CL) imaging of a select suite of grains indicates that they are internally characterized by concentric growth zonation consistent with an igneous origin. However, some grains display concentrically-zoned cores that are surrounded by thin, non-luminescent overgrowths interpreted to represent a second period of zircon growth (Fig. 3.9C inset). Three fractions without

143 significant visible overgrowths (Z1, Z2, and Z5) define a line with intercepts of

1679±2 Ma and 63±3 Ma (Fig. 3.9C). The upper intercept of 1679±2 Ma is interpreted to represent crystallization of the granitic sill, and the lower intercept of 63±3 Ma reflects more recent Pb loss perhaps related to Laramide tectonism. Two additional fractions (Z3 and Z4), one of which exhibits clear CL evidence for secondary growth (Z4, Fig. 3.9C inset), plot on top of one another, are slightly discordant (0.8 – 1.5%), and have 207Pb/206Pb ages of 1640 Ma and 1634 Ma, respectively (Table 3.1). A natural regression of these two points with the lower intercept pinned at 63 Ma (lower intercept of fractions Z1, Z2, and Z5; see above) yields an upper intercept age of 1640±3 Ma (Fig. 3.9C inset). This age is broadly consistent with metamorphic zircon ages from host-rock felsic gneiss (J01-RG2) and is interpreted to represent post-crystallization metamorphic zircon growth in the granite sill.

Southern Wet Mountains – Bear Creek/Williams Creek

The southern part of the range was chosen for new geochronology to determine the protolith ages of basement gneisses and the timing of metamorphism of gneisses and/or granites and development of penetrative, shallowly-dipping fabrics. Geochronologic studies of Bickford et al. (1989a) suggested that there were numerous Mesoproterozoic (ca. 1.4 Ga) intrusive phases that could be used to constrain fabric development in host rocks and in the granites themselves. The remarkably detailed mapping and descriptions of Boyer (1962) provided a template for more concentrated mapping and permitted the

144 correlation of local observations with surrounding areas. Additional mapping by Callahan (2002) and Perkins (2002) helped to establish the cross-cutting relationships described below. All of the samples described in the following section were collected from a single large exposure of gneiss and granite along the southwestern edge of the Wet Mountains near the intersection of Bear Creek and Williams Creek. This exposure, informally named “The Wall” because of its nearly 100 meter high vertical face (Fig. 3.4B), is 0.5 kilometers east of the Pole Creek Trailhead along San Isabel National Forest Road 630 (Fig. 3.6).

AMPHIBOLITE (J01-WC1). This sample of medium- to coarse-grained amphibolite comprises the dominant local host rock to the two generations of granitic intrusions (G2 and G3) described below. The sample contains a well- developed foliation that strikes E – W and dips moderately N and a moderately

NNW-plunging amphibole mineral lineation (Fig. 3.6A & B). Amphibolite host rock was sampled to determine the protolith age for basement gneisses and/or the timing of metamorphism across the southern part of the range. The sample yielded a relatively simple population of pink to tan, subrounded to subhedral, equant zircon interpreted to be metamorphic in origin. Three fractions (Z1 – Z3) plot close to concordia and have an average 207Pb/206Pb age of 1436±2 Ma (Fig.

3.10A). This age is interpreted to represent the timing of metamorphic zircon growth and/or recrystallization in the southern Wet Mountains contemporaneous

145

Figure 3.10. U-Pb concordia diagrams for samples from the Bear Creek/Williams Creek area, southern Wet Mountains. Ages are determined by linear regression through the data except where indicated, and probability of fit (%) is indicated in parentheses. See text for details.

146 with the emplacement of a widespread suite of coarse-grained granite sills (J01- WC2, see below).

COARSE-GRAINED GRANITE (J01-WC2). This sample was collected from one of numerous 1 – 5 m thick sills of coarse-grained to K-feldspar-megacrystic granite (G2 of Siddoway et al., 2000) exposed throughout the southern Wet Mountains. The sills contain a well-developed to locally-gneissic foliation (Fig. 3.11A) defined by biotite and dynamically-recrystallized feldspar and quartz that strikes ENE – WSW and dips moderately NNW (Fig. 3.6C). A pronounced mineral lineation defined by biotite and quartz plunges moderately down dip to the NNW, and asymmetric folds and mineral fabrics record reverse-sense (top-up-to-the- SSE) kinematics (Fig. 3.11B). This sample was collected to determine the timing of magmatism and gneissic, penetrative deformation across the southern part of the range. The sample yielded a simple population of pink to clear, euhedral to subhedral, prismatic zircon consistent with an igneous origin. Three fractions (Z1 – Z3) overlap concordia with an average 207Pb/206Pb age of 1435±4 Ma (Fig. 3.10B). This age is interpreted to represent emplacement and crystallization of the coarse-grained granitic sills and coincides with a regionally extensive suite of granitic plutons emplaced across the southern Rocky Mountains between 1440 – 1430 Ma (Reed et al., 1993). This age also agrees within error with the emplacement age of the Oak Creek pluton (1439±8 Ma; Bickford et al., 1989a), a

147 coarse-grained, deformed granitoid 40 km to the north in the Wet Mountains (Fig. 3.2). The granite also yielded abundant dark brown, angular titanite fragments. In thin section, titanite occurs parallel with the dominant gneissic fabric defined by dynamically-recrystallized feldspar, quartz, and biotite. This textural relationship suggests that titanite grew and/or recrystallized during metamorphism accompanying solid-state deformation of the granite sill. However, three fractions (T2, T4, and T5) define a line with intercepts of 1375±2 Ma and 43±65 Ma, and two fractions (T1 and T3) are colinear along a reference chord with intercepts of ca. 1362 Ma and 0 Ma (Fig. 3.10B). These five fractions are interpreted to represent two age populations, and the upper intercepts are interpreted to reflect closure ages that were thermally reset during emplacement of the San Isabel granite 5 – 10 km to the NW (Fig. 3.2). Both titanite ages agree with the published range of U-Pb zircon ages (1362±7 Ma and 1371±14 Ma, Bickford et al., 1989a) for the San Isabel granite. The presence of sapphirine in roof pendants of the nearby pluton suggests that rocks across the southern Wet Mountains experienced temperatures greater than 700°C during emplacement (Raymond et al., 1980). These temperatures are thought to be sufficient for resetting U-Pb systematics in titanite (Pidgeon et al., 1996; Verts et al., 1996).

FINE-GRAINED, FOLIATED GRANITE SILL (J01-WC3). This sample was collected from a fine-grained granite sill (G3 of Siddoway et al., 2000) that cuts both amphibolite and gneissic host-rocks and coarse-grained, gneissic G2 granite sills

148 (Fig. 3.11C). Sills range in thickness from 0.5 – 5.0 m and contain a well- developed solid-state biotite foliation that strikes ENE and dips moderately NNW (Fig. 3.6C). G3 sills also contain a moderately NNW-plunging biotite mineral lineation (Fig. 3.6D). Asymmetric mineral fabrics record reverse-sense (top-up- to-the-SSE) kinematics, and entire sills are locally deformed by 0.5 – 1.0 m wavelength asymmetric folds with E – W trending axes and southward vergence (Fig. 3.11D). This sample was collected to constrain the timing of fine-grained (G3) granitic magmatism and subsequent deformation across the southern part of the Wet Mountains. The granite yielded a single population of brown to tan, equant, euhedral to subhedral, prismatic zircon. Optical examination of the grains revealed slightly darker, translucent xenocrystic cores surrounded by lighter, clear, euhedral tips interpreted to represent igneous overgrowths. The small average grain size of the zircon population made it impossible to mechanically separate the core and rim material for analysis, and a shorter duration of air abrasion was required to preserve the volumetrically minor amount of overgrowth material. Four zircon fractions (Z1 – Z4) define a line with intercepts of 1749±28 Ma and 1390±10 Ma (Fig. 3.10C). We interpret the upper intercept to reflect zircon inheritance, and this age indirectly constrains the protolith age of basement gneisses in the southern Wet Mountains. The lower intercept is interpreted to represent the age of crystallization of the granitic sill, consistent with cross-cutting relationships described earlier.

149

Figure 3.11. Field photographs from the Bear Creek/Williams Creek area, southern Wet Mountains. A) View down on G2 granite float block with gneissic fabric defined by recrystallized K-feldspar and biotite. B) View to ENE of G2 granite in outcrop intruding amphibolite. Granite contains asymmetric (C-S) fabric defined by biotite that indicates reverse-sense (top-up-to-SSE) kinematics. Asymmetric fold in amphibolite just above sill also has southward vergence, suggesting reverse-sense kinematics.

150

Figure 3.11 continued. C) View down on G3 fine-grained granite sill subtly cutting across gneissic fabric (indicated by lines) in coarse-grained G2 granite. The G3 sill contains a fine-grained biotite foliation that is subparallel with G2 fabric. D) View to ENE of G3 granite sill containing well-developed foliation and small, asymmetric open folds. Folds have southward vergence that is consistent with reverse-sense (top-up-to-the-SSE) kinematic indicators elsewhere throughout the sill.

151 DISCUSSION

New structural observations from across the Wet Mountains coupled with U-Pb geochronology of basement gneisses and cross-cutting granitic rocks provide new constraints on the extent, timing, and style of Mesoproterozoic (ca. 1.4 Ga) deformation, metamorphism, and granitic magmatism in the southern Rocky Mountains (Table 3.2). These results indicate that the Wet Mountains record a long-lived but locally episodic history of ca. 1.4 Ga deformation and metamorphism that is both spatially and temporally associated with voluminous granitic magmatism and high-temperature metamorphism. These results further reveal that contrasting styles of deformation across the range were essentially coeval at ca. 1430 Ma, and subhorizontal penetrative deformation in the southern Wet Mountains persisted intermittently for 70 m.y. until emplacement of the ca. 1360 Ma San Isabel granite. The kinematics of deformation are broadly kinematically compatible across the range and are interpreted to reflect contrasting responses to subhorizontal, NNW-directed crustal shortening. Departures in kinematics or fabric orientations at shallower crustal levels are interpreted to represent the local influence of pre-existing crustal anisotropies (e.g., intrusion margins). The synchroneity of contrasting deformation styles indicates that a structural discontinuity in the middle crust was developed by ca. 1430 Ma and has important implications for tectonic models for ca. 1.4 Ga intracontinental between deformation and magmatism suggests that the two processes might have been fundamentally linked. A discussion of new results and

152

153

154 implications for both tectonic models and processes is presented in the following section.

Timing of ca. 1.4 Ga deformation, magmatism, and metamorphism across the Wet Mountains

Exposures in the northern Wet Mountains are characterized by discrete structural domains exhibiting dominantly subvertical folds and fabrics that record varying responses to ca. 1.4 Ga subhorizontal, NNW-directed shortening. The most intense deformation was localized along the 2 – 5 kilometer thick FPSZ, a subvertical, NNW-striking shear zone that might have accommodated multiple episodes of generally E-side-up, reverse-sense displacement. New geochronology indicates that at least one episode of shear zone displacement was accompanied by emplacement and deformation of a voluminous suite of coeval pegmatite dikes at 1431+5/-4 Ma. Syn-kinematic minerals occurring across the shear zone suggests that deformation was accompanied by metamorphism at temperatures exceeding 700°C (Givot, 1998). Siddoway et al. (2000) interpreted the geometry of the shear zone to represent part of a conjugate shear system developed during regional NW-directed shortening, but the shear zone geometry might also reflect more localized stress orientations that were controlled by the geometry of the large (>100 km2), relatively massive Crampton Mountain batholith adjacent to the

FPSZ. The gradual transition from the N-striking shear zone fabric (Ssz) to the NE-striking fabric of the Parkdale domain closely mirrors the mapped geometry of the southern margin of the batholith and, in part, supports the latter interpretation. High-temperature fabrics defined by dynamically-recrystallized

155 feldspar (plagioclase and K-feldspar), quartz, and biotite occur throughout Paleoproterozoic intrusive rocks dominating exposures along the eastern ARG (Parkdale domain), and local kinematic indicators are consistent with reverse- sense (NW-side-up-to-the-S) displacement. In the eastern ARG, new U-Pb titanite ages indicate that some degree of fabric development and/or reactivation occurred between 1422 – 1404 Ma. In the western ARG (Texas Creek domain),

E-trending, megascopic F2 folds formed at ca. 1.4 Ga after crystallization of ca. 1474 – 1450 Ma undeformed granitic plutons (Bickford et al., 1989a). Furthermore, ca. 1.43 Ga monazite inclusions that occur within cordierite in the limbs of F2 folds (Siddoway et al., 2002) indicate that metamorphism and folding might have been broadly synchronous with deformation along the adjacent FPSZ and in the Parkdale domain further to the east. Exposures in the central Wet Mountains are characterized by fabrics that strike E – W and have moderate to shallow NNW dips. These fabrics are both widespread and penetrative, affecting both basement gneisses and intrusive granitic rocks. Lanzirotti (1988) found evidence for three episodes of penetrative deformation, all attributed to NW – SE shortening, in the western part of the central Wet Mountains (Mt. Tyndall quadrangle; Brock and Singewald, 1968).

The first two deformation events (D1 and D2) are broadly correlated with ca. 1.7 Ga late-tectonic plutons exposed in the northern Wet Mountains and are constrained locally by a 1692±5 Ma U-Pb zircon age (Bickford et al., 1989a) from granulites in the central part of the range (Lanzirotti, 1988; Brock and Singewald,

1968). Lanzirotti (1988) suggested that the third deformation (D3) might have

156 coincided with the emplacement of a suite of younger deformed plutons (1650 –

1615 Ma; Bickford et al., 1989a), but D3 might have been contemporaneous with ca. 1.4 Ga metamorphism and magmatism. New U-Pb zircon and titanite data from 10 km to the east (Rattlesnake Gulch, Fig. 3.5) suggest that amphibolite and felsic gneiss formed between 1691 – 1680 Ma and experienced at least one episode of deformation involving foliation development and isoclinal folding prior to emplacement of a fine-grained granite sill at 1679±2 Ma. Whereas the felsic gneiss and sill record metamorphic zircon growth and/or recrystallization at ca. 1650 Ma, amphibolite experienced high-temperature metamorphism between 1436 – 1390 Ma likely accompanied by at least one episode of deformation involving fabric overprinting and/or reactivation. The youngest concordant zircon from amphibolite indicates that metamorphism occurred at ca. 1436 Ma, contemporaneous with metamorphism and emplacement of voluminous, coarse- grained G2 granite to the south. Metamorphic recrystallization and/or growth of titanite in the amphibolite accompanied local fabric reactivation at ca. 1390 Ma and coincided with the emplacement of a suite of fine-grained G3 granites across the southern part of the range.

New results from the southern Wet Mountains reveal that the earliest recorded phase of high-temperature, penetrative deformation occurred during emplacement of coarse-grained (G2) granitic sills at 1435±4 Ma. Magmatism was accompanied by foliation development (average orientation = 283, 56°N) and extensive recrystallization and/or new growth of metamorphic zircon in host-rock amphibolite at 1436±2 Ma. The observation that granitic sills locally cut a pre-

157 existing host-rock foliation requires the existence of at least one phase of earlier, likely Paleoproterozoic, deformation and metamorphism. However, any isotopic record of these events has effectively been overprinted by overwhelming ca. 1.4 Ga thermal effects. Deformation recorded by gneissic G2 granite sills indicates

NNW – SSE crustal shortening (average foliation = 243, 43°N) accompanied by reverse-sense (top-up-to-SSE) shearing following emplacement (average lineation

= 41°/243). The gneissic fabric in both G2 granites and host rocks is cut by 1390±10 Ma fine-grained (G3) granite sills, and these sills, in turn, record another pulse of NNW – SSE directed shortening and reverse-sense (top-up-to-SSE) shearing. All of the fabrics in the southern Wet Mountains are cut by the 1362±7 Ma San Isabel granite (Bickford et al., 1989). Although some exposures of San Isabel granite display evidence for complex local magmatic to solid-state deformation, it is largely (>90%) undeformed and, thus, provides a minimum age for all ca. 1.4 Ga deformation in the southern part of the range.

Contrasting levels of exposure from N – S across the Wet Mountains

The data and observations described above indicate that contrasting structural styles in the northern and southern Wet Mountains evolved at the same time between ca. 1436 – 1360 Ma and coincided with a widespread regional A- type granitic magmatism (Reed et al., 1993). Subvertical, strongly localized deformation in the northern part of the range is consistent with other documented regional examples of ca. 1.4 Ga deformation (e.g., Jessup et al., 2005; Shaw et al., 2001), and the FPSZ is another example of a crustal-scale shear zone that was

158 active during Mesoproterozoic intracontinental tectonism. In contrast, shallowly- to moderately-dipping, penetrative deformation exhibited to the south across the range is somewhat unique in the Mesoproterozoic record of the southern Rocky Mountains. The contrast in the style of contemporaneous deformation from N – S across the Wet Mountains is interpreted to reflect deeper levels of crustal exposure in the central and southern parts of the range relative to the ARG (Fig. 3.12). This interpretation is supported by regional 40Ar/39Ar thermochronology, contrasting styles and occurrence of ca. 1.4 Ga magmatism, and geobarometry (e.g., Shaw et al., 2005; Siddoway et al., 2000; Cullers et al., 1993). The central and southern parts of the Wet Mountains lie within the northern part of a region comprising southern Colorado and northern New Mexico that is characterized by a narrow range of Mesoproterozoic (ca. 1.45 – 1.35 Ga) hornblende and mica 40Ar/39Ar ages (Shaw et al., 2005). These ages contrast with other parts of northern and central Colorado that record similar mica ages but yield older hornblende ages ranging from ca. 1.7 – 1.4 Ga. Shaw et al. (2005) interpreted the region of consistently young hornblende and mica ages to come from deeper levels of crust at 1.4 Ga temperatures exceeding 500ºC. Surrounding areas with more variable hornblende ages are interpreted to represent shallower crustal levels that experienced temperatures between 300 – 500ºC at ca. 1.4 Ga (Shaw et al., 1999, 2005). Published estimates of emplacement depths of ca. 1.4 Ga plutons across the region typically range from 7 – 15 km (200 – 500 MPa; Williams, 1991; Cullers et al., 1993; Barinek et al., 1999). In contrast, the 1362 Ma San Isabel granite (Bickford et al., 1989a) in the southern Wet Mountains was

159 emplaced at depths of 17 – 23 km (500 – 700 Mpa; Cullers et al., 1992) based on geobarometry (Al-in-hornblende) and the presence of primary, euhedral magmatic epidote. Whereas the region yielding young hornblende and mica ages is also characterized by a relative paucity of discrete ca. 1.4 Ga granitic plutons, new results described above reveal that the southern Wet Mountains experienced two episodes of granitic magmatism at 1430 Ma and 1390 Ma, respectively, prior to emplacement of the 1362 Ma San Isabel pluton. The G2 and G3 granites volumetrically represent batholith-sized intrusions but were emplaced as a concordant magmatic “framework” rather than as discrete plutonic bodies. The change in style of magmatism coincides with the change in structural style from N – S across the range, and concordant granites are interpreted to represent deeper- level equivalents, and perhaps magmatic feeders, of contemporaneous, shallower mid-crustal plutons (Fig. 3.12).

Ca. 1.4 Ga mid- to lower-crustal flow

Although fabrics developed during penetrative deformation in the central and southern part of the range presently dip moderately (50º) to the north, evidence for post-1.4 Ga differential uplift and tilting of the crustal section suggests that fabrics dipped more shallowly during their formation. Contrasting levels of exposure indicated by mineral assemblages and geobarometry between the northern and southern Wet Mountains require differential exhumation that would produce a minimum of 10° – 15° of crustal tilt. Furthermore, apatite- fission-track thermochronology from N – S across the Wet Mountains indicates

160 much younger differential uplift of a Late Cretaceous apatite partial annealing zone and the Eocene Rocky Mountain erosion surface (Kelley and Chapin, 2004). Restoration of the contrasting surface elevations across the range suggests that an additional >10° of crustal tilt occurred during the Cenozoic. Thus, the restoration of ca. 1.4 Ga fabrics across the southern Wet Mountains to shallower dips reflecting subhorizontal crustal flow is considered to be reasonable. Following local and regional arguments that the southern Wet Mountains represent relatively deep levels of crustal exposure, ca. 1.4 Ga penetrative deformation across this part of the range is interpreted to represent shallowly-dipping to subhorizontal crustal flow. The earliest recognized episode of penetrative flow accompanied the emplacement of coarse-grained G2 granites at 1435 Ma and continued, perhaps intermittently, ca. 70 m.y. until emplacement of the San Isabel granite. The orientation and kinematics of deformation remained remarkably consistent during this time and are consistent with long-lived, subhorizontal NNW – SSE crustal shortening.

Implications for Mesoproterozoic tectonic models

Reconciling the contrasting, yet coeval, styles of ca. 1.4 Ga deformation across the Wet Mountains requires a structural discontinuity in the middle crust that evolved by ca. 1430 Ma and persisted until ca. 1360 Ma. Shaw et al. (2005) proposed a model in which the discontinuity corresponds with large volumes of

161

Figure 3.12. Schematic block diagram of the middle crust beneath southern Colorado at ca. 1.4 Ga. Interpreted, reconstructed relationships illustrate contrasting structural and magmatic styles with relative position (i.e., depth) in the crust. Shallower levels of exposure are characterized by steeply dipping to subvertical fabrics, and deformation is commonly localized along discrete shear zones. Examples include the Five Points Shear Zone (FPSZ, Plate 2) and the Homestake Shear Zone (Shaw et al., 2001). Deeper levels of exposure are characterized by moderately to shallowly dipping, penetrative gneissic fabrics and penetrative deformation (southern Wet Mountains).

162 A-type granitic magma that were emplaced near the brittle-ductile transition where they would have encountered a rheologic and thermal barrier to further ascent (Fig. 3.12). This mid-crustal magmatic layer would have effectively insulated the deeper crust and maintained higher temperatures throughout the duration of magmatism. Higher temperatures throughout deeper levels of exposure are indicated by widespread <1.45 Ga hornblende 40Ar/39Ar ages across southern Colorado and northern New Mexico (Shaw et al., 2005). The influx of heat and magma from the lower crust significantly weakened the crust beneath the brittle-ductile transition and permitted long-lived penetrative deformation and flow throughout the lower crust. Rocks above the mid-crustal magma layer experienced widespread high-T, low-P metamorphism (Williams and Karlstrom, 1996), and argon systematics were completely reset in mica but only partially reset in hornblende (Shaw et al., 1999). Deformation at shallower crustal levels was localized along pre-existing crustal weaknesses (e.g., shear zones; Shaw et al., 2001; McCoy et al., 2005) or concentrated in thermally softened aureoles surrounding shallower ca. 1.4 Ga plutons (Aleinikoff et al., 1993). In some cases, the ascent of ca. 1.4 Ga plutons was locally accommodated by pull-apart structures in coeval transpressional shear systems (McCoy et al., 2005) and in areas where the dominant basement fabric was oriented at high angles to the regional shortening direction (e.g., 1434 Ma Music Pass pluton, Chapter 1 of this study). The precise location and geometry of the structural discontinuity and mid- crustal magma layer is difficult to identify in the Wet Mountains due to

163 incomplete exposure and local access problems but likely occurs south of the ARG. Exposures south of the ARG are characterized by an increasing abundance of granitic intrusions, and gneissic rocks show evidence for more complex, migmatitic behavior within 10 km south of the ARG (Collins et al., 2004). Despite these local complexities, the mechanical behavior of rocks above and below the discontinuity and the timing of deformation and magmatism in the northern and southern Wet Mountains, respectively, are consistent with the model proposed by Shaw et al. (2005) and represent an important new contribution to understanding ca. 1.4 Ga tectonism across the southern Rocky Mountains and southwestern U.S. Evidence for long-lived penetrative deformation that consistently records subhorizontal NNW – SSE shortening between 1435 – 1360 Ma supports other regional structural arguments for crustal shortening accompanying ca. 1.4 Ga granitic magmatism (e.g., Nyman et al., 1994). Furthermore, ca. 1.4 Ga granitic rocks across the central and southern Wet Mountains likely represent some of the most sensitive strain recorders across the region. The consistent style and orientation of deformation across such a large area (>200 km2) over ca. 70 m.y. is kinematically compatible with an abundance of deformation at shallower levels during the same time. The regionally consistent kinematics of the intracontinental system implies an organized, larger- scale deformational system that, in turn, is consistent with a convergent margin system, perhaps with an oblique character, along the distal southern margin of Laurentia (Nyman et al., 1994). Although direct evidence of a margin has not been identified, geochemical and isotopic data suggest that juvenile crust was

164 formed and added to Laurentia between 1.5 – 1.3 Ga (Bennett and DePaolo, 1987; Patchett and Ruiz, 1989). Models for intraplate deformation indicate that the strength of the lower crust fundamentally influences the geometry and strain distribution within an active orogen. Royden (1996) demonstrated that a weak lower crust underlying strong upper crust permits transmission of compressive stresses >1000 km from the site of convergence. This can produce a broad orogen with high average elevation but low relative relief. There is new evidence in the Wet Mountains for a structural discontinuity in the middle crust that developed at ca. 1430 Ma that was likely focused across the magma-rich layer emplaced along the brittle-ductile transition (Shaw et al., 2005). Penetrative deformation throughout the southern Wet Mountains is interpreted to result from lower crustal flow, and the close temporal relationship between deformation and magmatism suggests that ca. 1435 Ma granites played a fundamental role in weakening the lower crust and permitting the propagation of deformation across such a widespread region. Shallowly-dipping, penetrative fabrics combined with the voluminous, concordant framework of ca. 1.4 Ga granitic magmas might also represent a deeply-exhumed analog to low-viscosity layers that have been inferred in modern intracontinental orogenic settings like Tibet and the Altiplano (e.g., Nelson and Project INDEPTH, 1996; Brasse et al., 2002). Although granites played a fundamental role in localizing deformation at shallower crustal levels and in accommodating penetrative deformation at deeper levels, it is not clear whether or not there is a specific genetic link between ca. 1.4 Ga magmatism and tectonism throughout the southern Rocky Mountains. A-type

165 granites are not commonly associated with convergent tectonic environments and are not known to occur within modern intracontinental orogenic settings. Shaw et al. (2005) speculated that topographically driven syncontractional extension at shallower crustal levels at ca. 1.4 Ga could reconcile regional structural evidence for NW – SE shortening with the geochemistry of coeval magmatism. Modern orogens and orogenic plateaus commonly display evidence for synchronous shortening and extension both parallel and perpendicular to the orogen (e.g., Burchfiel et al., 1992). High heat flow required for generating A-type granites through melting of the lower crust could be related to asthenospheric upwelling driven by thermal or convective removal of an orogenically thickened lithosphere (e.g., Houseman et al., 1981) or by emplacement of a large mantle plume (Hoffman, 1989; Ferguson et al., 2004). The former mechanism is preferred because it is a process believed to occur quite commonly in convergent orogenic systems (e.g., Collins, 1984; Cloos, 2003) and has been inferred along the eastern margin of Laurentia during approximately the same time (Rivers, 1997).

SUMMARY/CONCLUSIONS

New field mapping coupled with U-Pb zircon and titanite geochronology from the Wet Mountains, Colorado, demonstrate that contrasting styles of deformation from N – S across the range were coeval at ca. 1430 Ma and accompanied a widespread regional pulse of A-type granitic magmatism. Deformation in the northern part of the range is characterized by subvertical fabrics and folding and involved localization of strain among three structural

166 domains. At least one phase of movement along the Five Points shear zone was accompanied by pegmatite diking at 1430+5/-3 Ma, and kinematic indicators across the zone indicate that deformation involved reverse-oblique, E-side-up displacement. Although timing constraints are less well constrained, ca. 1.4 Ga deformation in adjacent structural domains included upright, megascopic folding and reverse-sense, high-temperature reactivation of NE-striking tectonic fabrics. Deformation throughout the central and southern Wet Mountains was penetrative in character and produced multiple generations of moderately- to shallowly- dipping fabrics and isoclinal folds. The earliest pulse of Mesoproterozoic deformation and high-temperature metamorphism accompanied the emplacement of a concordant network of coarse-grained granite sills at 1435±4 Ma. These sills contain a concordant, gneissic foliation that is cut by a younger suite of fine- grained granite sills that were emplaced at 1390±10 Ma and contain a concordant biotite foliation. Both generations of granite as well as surrounding host gneisses have fabric orientations that suggest NNW – SSE crustal shortening and asymmetric folds and mineral fabrics that consistently indicate reverse-sense (top- up-to-SSE) displacement. Thermochronologic and geobarometric data coupled with structural observations suggest that the southern Wet Mountains represent some of the deepest levels of Proterozoic crustal exposure throughout the Rocky Mountains (Shaw et al., 2005; Cullers et al., 1993), and long-lived (1435 – 1360 Ma), shallowly-dipping penetrative deformation across the central and southern part of the range is interpreted to represent subhorizontal crustal flow facilitated, and

167 perhaps accommodated, by a voluminous framework of coeval granite. Evidence for subhorizontal NNW – SSE crustal shortening across the range is consistent with tectonic models suggesting that ca. 1.4 Ga granites were emplaced during regional convergent orogenesis perhaps within an intracontinental orogenic plateau. Penetrative lower crustal flow and ca. 1430 Ma structural decoupling between shallower and deeper crustal levels indicate that ca. 1435 Ma magmatism fundamentally weakened the lower crust and allowed deformation to propagate far inboard (1000 km) from the locus of active convergence. Furthermore, shallowly- to moderately-dipping, melt-dominated deformation at lower crustal levels is interpreted to represent a deeply-exhumed analog for low-viscosity layers inferred beneath modern intracontinental orogenic systems.

168 Appendix 1

METHODS: ISOTOPE DILUTION THERMAL IONIZATION MASS SPECTROMETER (ID-TIMS) ANALYSIS

All sample processing was done at The University of Texas at Austin. Rock samples were crushed to mineral size under clean conditions by using a jaw crusher and disc pulverizer. Initial mineral separation was done using a Wilfley table, and heavy-mineral components were processed further with sieves, heavy liquids, and a Frantz magnetic separator. Mineral fractions were characterized by using a binocular reflected-light microscope, a transmitted-light petrographic microscope (with condenser lens inserted to minimize edge refraction), and a scanning cathodoluminescence (CL) imaging system on a JEOL T330A scanning electron microscope (SEM). Multiple or single grains of each population were selected for analysis on the basis of optical properties to ensure that only the highest-quality grains were analyzed. All mineral fractions analyzed were strongly abraded (Krogh, 1982), were subsequently reevaluated optically, and then were washed successively in distilled 4N nitric acid, water, and acetone. They were loaded dry in to Teflon capsules with a mixed 205Pb-235U isotopic tracer solution and dissolved with appropriate acids (HF and HNO3 for silicates, 6.2N HCl for monazite). Chemical separation of U and Pb from zircon using minicolumns (0.044 mL resin volume; after Krogh, 1973) resulted in a total Pb procedural blank of 1 – 2 pg over the period of analyses. Chemical separation from titanite on 0.250 mL columns resulted in a total procedural Pb blank of 1 – 2 pg. The U procedural blank is

169 estimated to have been 0.5 pg for both column types. Pb and U were loaded together with silica gel and phosphoric acid onto an outgassed filament of zone- refined rhenium ribbon and analyzed on a multicollector MAT 261 thermal- ionization mass spectrometer, either operating in static mode (with 204Pb in the axial secondary-electron multiplier [SEM]-ion-counting system) or dynamic mode with all masses measured sequentially by the SEM-ion-counting system. Ages were calculated by using decay constants of Jaffey et al. (1971). Errors on isotopic ratios were calculated by propagating uncertainties in measurement of isotopic ratios, fractionation, and amount of blank with a program modified after algorithms by L. Heaman (University of Alberta, Edmonton). Results are reported in Tables 1.1, 2.1, and 3.1 with 2σ errors. Linear regressions were performed with the procedure of Davis (1982). The goodness of fit of a regressed line is represented as a probability of fit, where 10% or better is considered acceptable and corresponds to a mean square of weighted deviates (MSWD) of 2 or less. Ages listed in the text, tables, and figures are given with 2σ errors.

170 Appendix 2

METHODS: LASER ABLATION INDUCTIVELY COUPLED MASS SPECTROMETER (LA-ICP-MS) ANALYSIS

Samples were processed and zircons were separated and hand-picked for LA-ICP-MS analysis by the same methods as outlined for TIMS analysis (Appendix 1). Selected zircons of comparable size were placed on two-sided tape, collared, and covered with epoxy. The zircons in the resulting puck were ground and polished to approximately two thirds of their original thickness. Cathodoluminescent images were made to characterise the zircons and guide subsequent analysis. Laser ablation analysis utilised a Merchantek 213nm YAG- laser connected to a Micromass quadrupole mass spectrometer (Platform). Fractionation and inherent detector non-linearity were accounted for by analysing zircons already well characterised by TIMS. Corrections necessary to obtain the correct 207Pb/206Pb ratios for these internal laboratory standards, covering a range of intensities and isotopic ratios, were applied to unknowns. Standards were run throughout each session. Blanks and off-peak baselines were also determined throughout each analytical session. Selecting and analysing only the highest quality zircons precluded the need for common Pb corrections, a procedure made difficult by high 204Hg counts due to low-level contaminates in the argon gas. A single zircon analysis comprises 493 ten microsecond scans of 207Pb and 206Pb. Ratios reflecting the moving average of 20 207Pb and 206Pb measurements are first plotted on a graph to check for anomalous ratios throughout a run. Despite a pre-ablation pass to clean the sample surface, some 171 grains exhibited high ratios at the beginning of an analysis: these scans were presumed to reflect common Pb and were removed from further consideration. A jump from one ratio plateau to another during one analysis is interpreted to reflect piercing a core of different age. Only data from one plateau at a time are considered. In cases where the beam pierces the grain too deeply and ejecta are not effectively emitted towards the end of an analysis, the signal intensity and commonly also isotope ratios change dramatically. Data from the latter part of such runs are also rejected. With this first assessment of data complete, 207Pb and 206Pb were then passed through a 4-sigma filter to remove anomalous counts before being passed through a more rigorous 2-sigma filter. The averages of the remaining individual measurements (typically <5% rejection) of 207Pb and 206Pb provided the final 207Pb/206Pb ratio and consequent age. Given the transient signal inherent in LA- ICP-MS and sequential acquisition required by the single collector, 2 sigma errors reflect the standard deviation of a moving average of ratios based on 20 scans.

172 Appendix 3

QUARTZITE LA-ICP-MS DATA RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std

J01-BR1: Quartzite from Blue Ridge, Colorado J01-BR1 1 74005 9156 479426 418 0.12373 0.11653 1904 3.4 J01-BR1 2 33875 4025 165807 442 0.11881 0.11261 1842 6.8 J01-BR1 3 34587 3788 168595 472 0.10953 0.10353 1688 4.3 J01-BR1 4 36448 4399 164750 384 0.12070 0.11451 1872 4.2 J01-BR1 5 52917 6036 261814 475 0.11407 0.10798 1766 4.8 J01-BR1 6 62913 7187 333482 457 0.11423 0.10785 1763 3.3 J01-BR1 7 56266 6731 263669 454 0.11964 0.11330 1853 3.5 J01-BR1 8 69165 7769 341876 470 0.11233 0.10579 1728 3.1 J01-BR1 9 25804 3003 115456 475 0.11637 0.11010 1801 4.5 J01-BR1 10 40350 4477 274747 472 0.11095 0.10497 1714 4.6 J01-BR1 11 83793 9334 411777 478 0.11139 0.10430 1702 3.3 J01-BR1 12 42846 4841 204668 378 0.11298 0.10703 1749 4.6 J01-BR1 13 31735 3548 155596 355 0.11181 0.10573 1727 4.3 J01-BR1 14 15305 1760 74884 384 0.11500 0.10852 1775 4.8 J01-BR1 15 41978 4721 213488 477 0.11247 0.10652 1741 5.0 J01-BR1 19 49998 5584 270189 263 0.11169 0.10572 1727 3.5 J01-BR1 20 23194 2622 195307 218 0.11302 0.10681 1746 2.9 J01-BR1 21 105736 13171 468500 362 0.12457 0.11575 1892 3.5 J01-BR1 22 41569 4957 190554 325 0.11925 0.11318 1851 3.8 J01-BR1 23 23601 2800 105952 239 0.11863 0.11228 1837 5.6 J01-BR1 24 48394 5397 233767 469 0.11152 0.10557 1724 4.4 J01-BR1 25 52917 5742 265910 442 0.10851 0.10258 1671 4.1 J01-BR1 26 91802 10483 446784 474 0.11419 0.10660 1742 3.4 J01-BR1 27 53605 6223 261263 436 0.11609 0.10993 1798 4.3 J01-BR1 28 91814 10254 455454 473 0.11168 0.10422 1701 3.7 J01-BR1 29 22362 2508 109882 478 0.11216 0.10597 1731 5.3 J01-BR1 30 84444 9668 438365 473 0.11448 0.10721 1753 4.0 J01-BR1 31 109096 13149 513110 472 0.12052 0.11177 1828 3.2 J01-BR1 32 32825 3786 151653 469 0.11535 0.10921 1786 5.8 J01-BR1 33 41923 4772 215467 438 0.11383 0.10786 1764 4.9 J01-BR1 34 62931 7388 298418 475 0.11740 0.11091 1814 4.1 J01-BR1 35 41001 4856 203226 474 0.11845 0.11239 1838 4.9 J01-BR1 36 94274 10645 472219 474 0.11292 0.10528 1719 3.1 173 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J01-BR1 37 73549 8809 358908 406 0.11976 0.11275 1844 3.2 J01-BR1 38 80464 9318 405994 458 0.11581 0.10865 1777 2.8 J01-BR1 39 31302 3542 150169 464 0.11316 0.10704 1750 5.3 J01-BR1 40 120855 14228 581752 464 0.11773 0.10858 1776 3.0 J01-BR1 41 35346 3845 175430 300 0.10878 0.10123 1647 4.7 J01-BR1 42 37073 4326 168250 476 0.11669 0.10901 1783 5.3 J01-BR1 43 44932 5148 224356 255 0.11457 0.10729 1754 4.8 J01-BR1 44 56576 6321 283590 368 0.11173 0.10466 1708 4.5 J01-BR1 45 51477 5925 263586 473 0.11509 0.10791 1764 5.1 J01-BR1 46 53933 6141 335749 362 0.11387 0.10674 1744 4.9 J01-BR1 47 36891 4290 383428 289 0.11629 0.10862 1776 3.3 J01-BR1 48 84684 9813 442171 474 0.11587 0.10820 1769 3.5 J01-BR1 49 43271 5238 203207 333 0.12104 0.11358 1857 3.3 J01-BR1 50a 106228 14301 543306 286 0.13462 0.12544 2035 2.9 J01-BR1 50b 146568 23608 739203 175 0.16107 0.14947 2340 4.1 J01-BR1 51 41267 4674 204550 206 0.11327 0.10589 1730 5.0 J01-BR1 52 42839 4940 257150 241 0.11531 0.10796 1765 3.4 J01-BR1 53 28959 7486 74718 348 0.25852 0.24885 3177 5.0 J01-BR1 54 41286 4552 205550 476 0.11026 0.10295 1678 6.4 J01-BR1 55 64549 7447 307250 474 0.11537 0.10813 1768 6.1 J01-BR1 56 83952 9620 415609 234 0.11459 0.10699 1749 4.0 J01-BR1 57 69240 8089 353583 317 0.11683 0.10946 1790 5.7 J01-BR1 58 75354 8645 383219 330 0.11473 0.10732 1754 4.0 J01-BR1 59 57385 6465 264196 106 0.11266 0.10557 1724 5.0 J01-BR1 60 38737 4243 194211 475 0.10953 0.10209 1662 6.2 J01-BR1 61 53439 6301 266181 374 0.11792 0.11067 1810 3.9 J01-BR1 62 99075 11771 519913 285 0.11881 0.11059 1809 6.9 J01-BR1 63 58868 9728 183752 320 0.16525 0.15656 2419 3.9 J01-BR1 64 67145 7534 332966 410 0.11220 0.10504 1715 4.7 J01-BR1 65 53530 6362 268050 379 0.11885 0.11157 1825 4.2 J01-BR1 66 61935 7038 309468 343 0.11363 0.10648 1740 3.6 J01-BR1 67 29348 3276 139955 334 0.11162 0.10379 1693 5.0 J01-BR1 68 68774 7744 339593 427 0.11260 0.10540 1721 4.9 J01-BR1 69 38830 4592 190796 403 0.11826 0.11068 1811 6.4 J01-BR1 70 57649 6825 332632 473 0.11839 0.11112 1818 5.5 J01-BR1 71 42357 4959 199582 357 0.11707 0.10966 1794 7.2 J01-BR1 72 65861 8171 292115 469 0.12407 0.11651 1903 5.4 J01-BR1 73 54605 6213 264073 435 0.11379 0.10666 1743 4.3 J01-BR1 74 53194 10837 503825 325 0.20373 0.19424 2778 5.4

174 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J01-BR1 75 48435 5299 236893 331 0.10940 0.10233 1667 5.2 J01-BR1 76 43843 5500 273018 264 0.12545 0.11791 1925 5.0 J01-BR1 77 60897 6996 310333 380 0.11489 0.10770 1761 6.0 J01-BR1 78 29153 3350 134321 469 0.11493 0.10698 1749 7.1 J01-BR1 79 49023 5748 279842 415 0.11725 0.10999 1799 5.6 J01-BR1 80 71923 13240 223065 465 0.18408 0.17436 2600 6.1

J01-BR2: Basal conglomerate from Blue Ridge, Colorado J01-BR2 1 138751 15729 741422 470 0.11336 0.10369 1691 6.4 J01-BR2 2 45637 5086 247629 310 0.11144 0.10552 1723 4.2 J01-BR2 3 37774 4172 194798 289 0.11045 0.10446 1705 4.3 J01-BR2 4 46157 5068 237725 289 0.10980 0.10391 1695 4.4 J01-BR2 5 62078 6758 321553 379 0.10886 0.10269 1673 3.8 J01-BR2 6 56622 6265 310142 385 0.11064 0.10456 1707 3.0 J01-BR2 7 89024 10095 478926 465 0.11340 0.10598 1731 4.2 J01-BR2 8 271296 30408 1692729 475 0.11208 0.09758 1578 2.8 J01-BR2 9 68191 7706 367481 333 0.11300 0.10647 1740 2.7 J01-BR2 10 172222 20389 893514 469 0.11839 0.10709 1750 4.2 J01-BR2 11 53306 5885 269457 404 0.11040 0.10440 1704 3.7 J01-BR2 12 57831 6711 331089 377 0.11604 0.10977 1796 3.2 J01-BR2 13 47150 5292 239951 474 0.11224 0.10629 1737 4.0 J01-BR2 14 58410 6347 288890 472 0.10867 0.10260 1672 4.1 J01-BR2 15 56821 6091 307696 357 0.10719 0.10122 1647 3.7 J01-BR2 16 50737 5635 254788 329 0.11107 0.10511 1716 4.9 J01-BR2 17 52110 5811 260537 475 0.11151 0.10551 1723 5.7 J01-BR2 18 32165 3642 168126 232 0.11323 0.10712 1751 4.8 J01-BR2 19 42915 4939 229550 467 0.11508 0.10909 1784 6.7 J01-BR2 20 53911 6364 334660 286 0.11804 0.11181 1829 4.1 J01-BR2 21 52052 5677 281131 191 0.10906 0.10312 1681 3.5 J01-BR2 22 80894 9035 407058 474 0.11169 0.10470 1709 3.4 J01-BR2 23 38049 4172 194067 472 0.10965 0.10368 1691 4.9 J01-BR2 24 150361 17092 797476 471 0.11367 0.10349 1688 2.7 J01-BR2 25 63032 7109 359324 263 0.11279 0.10645 1739 4.2 J01-BR2 26 71579 8142 387538 289 0.11375 0.10706 1750 3.6 J01-BR2 27 79500 8912 418455 471 0.11210 0.10515 1717 4.8 J01-BR2 28 45184 4881 225716 242 0.10803 0.10218 1664 5.3 J01-BR2 29 42186 4881 224733 191 0.11570 0.10970 1794 3.1 J01-BR2 30 54010 6195 268976 258 0.11471 0.10858 1776 2.9 J01-BR2 31 73319 8094 386775 240 0.11039 0.10377 1693 5.8

175 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J01-BR2 32 40567 4492 207266 193 0.11072 0.10475 1710 4.3 J01-BR2 33 44922 5001 227917 470 0.11134 0.10542 1722 4.2 J01-BR2 34 55292 6122 296119 379 0.11073 0.10468 1709 3.9 J01-BR2 35 38829 4357 195413 380 0.11222 0.10620 1735 4.6 J01-BR2 36 52379 5792 266044 210 0.11058 0.10460 1707 3.6 J01-BR2 37 46805 5172 243054 334 0.11051 0.10460 1707 3.2 J01-BR2 38 65889 7791 411247 474 0.11825 0.11161 1826 4.8 J01-BR2 39 87683 10049 457640 377 0.11461 0.10718 1752 3.4 J01-BR2 40 41550 4614 208774 447 0.11106 0.10374 1692 4.3 J01-BR2 41 54039 6137 311871 475 0.11356 0.10643 1739 6.2 J01-BR2 42 48829 5455 247509 283 0.11172 0.10460 1707 3.7 J01-BR2 43 47410 5379 251343 384 0.11347 0.10627 1736 6.1 J01-BR2 44 67084 7594 337583 469 0.11320 0.10600 1732 4.1 J01-BR2 45 42557 4741 234111 266 0.11140 0.10412 1699 3.7 J01-BR2 46 48227 5273 264144 477 0.10934 0.10226 1666 4.5 J01-BR2 47 50987 5633 257324 476 0.11047 0.10341 1686 4.1 J01-BR2 48 42539 4744 209134 333 0.11151 0.10423 1701 3.4 J01-BR2 49 51529 5833 288707 471 0.11321 0.10608 1733 4.8 J01-BR2 50 36546 4134 199946 293 0.11312 0.10551 1723 3.5 J01-BR2 51 53858 6100 301844 234 0.11327 0.10615 1734 2.9 J01-BR2 52 54789 6243 331648 213 0.11394 0.10681 1746 4.1 J01-BR2 53 40093 4442 242696 234 0.11078 0.10337 1685 5.0 J01-BR2 54 61437 6850 339520 355 0.11150 0.10442 1704 5.6 J01-BR2 55 53126 5968 296768 396 0.11233 0.10523 1718 5.2 J01-BR2 56 67241 7750 336339 447 0.11526 0.10799 1766 6.1 J01-BR2 57 64540 7042 335642 194 0.10912 0.10209 1662 4.9 J01-BR2 58 51918 5794 266095 440 0.11160 0.10452 1706 5.1 J01-BR2 59 82040 9383 430157 414 0.11437 0.10682 1746 4.6 J01-BR2 60 63030 7064 322692 471 0.11207 0.10495 1713 4.8 J01-BR2 61 65000 7169 322889 301 0.11028 0.10321 1683 5.2 J01-BR2 62 71336 8203 378245 437 0.11499 0.10765 1760 4.5 J01-BR2 63 65348 7378 383989 341 0.11290 0.10573 1727 5.5 J01-BR2 64 45264 5175 242913 379 0.11433 0.10707 1750 6.6 J01-BR2 65 40420 4526 201245 366 0.11197 0.10459 1707 4.6 J01-BR2 66 35928 3971 208283 233 0.11052 0.10295 1678 6.3 J01-BR2 67 65106 8110 399500 237 0.12457 0.11701 1911 5.1 J01-BR2 68 59455 6880 332724 337 0.11572 0.10852 1775 4.9 J01-BR2 69 58094 6291 289770 471 0.10829 0.10133 1649 6.1 J01-BR2 70 61312 6872 299472 328 0.11208 0.10498 1714 5.1

176 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J01-BR2 71 61831 7034 329237 470 0.11375 0.10660 1742 7.2 J01-BR2 72 186655 21530 1040161 383 0.11535 0.10473 1710 4.6 J01-BR2 73 40001 4455 207682 379 0.11136 0.10393 1695 3.4 J01-BR2 75 76009 8517 398711 229 0.11205 0.10473 1710 5.0 J01-BR2 76 65952 7498 346820 285 0.11368 0.10648 1740 3.5 J01-BR2 77 99671 11142 489669 464 0.11179 0.10389 1695 3.3 J01-BR2 78 54071 6061 282182 334 0.11209 0.10500 1714 4.0 J01-BR2 79 49983 5660 260408 365 0.11324 0.10609 1733 3.3 J01-BR2 80 63203 7029 337075 477 0.11121 0.10412 1699 3.9

J03-PC1: Quartzite from Phantom Canyon, Colorado J03-PC1 1 87371 9226 622432 248 0.10559 0.10206 1662 3.0 J03-PC1 2 89878 9670 583770 479 0.10759 0.10361 1690 3.1 J03-PC1 3 72797 8152 1687539 482 0.11199 0.10995 1799 3.6 J03-PC1 4 34262 3652 182274 472 0.10658 0.10494 1713 4.1 J03-PC1 5 70939 7643 499739 473 0.10774 0.10598 1731 3.5 J03-PC1 6 72629 7837 373273 472 0.10790 0.10598 1731 3.2 J03-PC1 7 44643 4894 455176 263 0.10962 0.10915 1785 4.0 J03-PC1 8 63523 6641 354203 336 0.10455 0.10343 1687 3.8 J03-PC1 9 53775 5132 819647 287 0.09543 0.09450 1518 3.3 J03-PC1 10 70709 7622 1690677 471 0.10780 0.10607 1733 3.9 J03-PC1 11 64326 6783 997182 477 0.10545 0.10427 1702 2.3 J03-PC1 12 51316 5365 401939 340 0.10455 0.10394 1696 2.9 J03-PC1 13 56652 6104 723084 477 0.10774 0.10706 1750 3.1 J03-PC1 14 115450 12463 610197 410 0.10795 0.10055 1634 2.2 J03-PC1 15 76427 8177 889910 474 0.10699 0.10469 1709 2.5 J03-PC1 16 57547 6278 1458158 243 0.10909 0.10838 1772 3.5 J03-PC1 17 86932 9139 607317 477 0.10512 0.10167 1655 3.1 J03-PC1 18 36654 3751 188809 362 0.10233 0.10080 1639 3.8 J03-PC1 19 31499 3253 161977 420 0.10329 0.10100 1643 4.0 J03-PC1 20 92982 9968 1186062 475 0.10720 0.10282 1676 2.8 J03-PC1 21 43374 4634 287132 335 0.10685 0.10620 1735 3.0 J03-PC1 22 95654 10324 699429 474 0.10793 0.10312 1681 2.8 J03-PC1 23 84170 9125 827704 477 0.10842 0.10513 1717 3.9 J03-PC1 24a 106205 12424 561063 195 0.11698 0.10994 1798 1.8 J03-PC1 24b 75332 8094 385674 285 0.10745 0.10525 1719 3.1 J03-PC1 25 66072 6914 359298 475 0.10464 0.10333 1685 2.6 J03-PC1 26 46746 4974 249494 332 0.10639 0.10584 1729 2.3 J03-PC1 27 149733 16467 818167 472 0.10998 0.10998 1799 3.7

177 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J03-PC1 28 103825 11518 762105 422 0.11094 0.10475 1710 2.3 J03-PC1 29 39674 5651 199815 147 0.14243 0.14323 2267 5.3 J03-PC1 30 24279 2644 119128 469 0.10890 0.10551 1723 4.9 J03-PC1 31 71642 7529 619927 472 0.10510 0.10331 1685 3.9 J03-PC1 32 79144 8545 453608 475 0.10797 0.10532 1720 3.7 J03-PC1 33 136997 15849 840285 317 0.11569 0.10499 1714 4.0 J03-PC1 34 46605 4779 354207 473 0.10255 0.10183 1658 4.3 J03-PC1 35 91753 9723 512226 467 0.10597 0.10184 1658 4.4 J03-PC1 36 51066 5358 274013 477 0.10493 0.10433 1703 3.9 J03-PC1 37 58099 6145 479824 474 0.10576 0.10497 1714 4.2 J03-PC1 38 50920 9061 174643 477 0.17794 0.17801 2634 4.4 J03-PC1 39 34828 3644 1075102 290 0.10462 0.10295 1678 3.7 J03-PC1 40 57904 5986 550344 340 0.10337 0.10255 1671 4.5 J03-PC1 41 70200 7491 383253 471 0.10671 0.10505 1715 3.2 J03-PC1 42 84504 8999 437135 474 0.10650 0.10327 1684 4.2 J03-PC1 43 112431 11536 1019346 238 0.10261 0.10261 1672 3.1 J03-PC1 44 120357 13432 1837646 478 0.11160 0.10326 1683 2.4 J03-PC1 45 45285 4670 339231 285 0.10312 0.10238 1668 3.9 J03-PC1 46 204356 22006 1083376 476 0.10768 0.10768 1761 3.3 J03-PC1 47 60471 6414 611440 479 0.10606 0.10515 1717 3.1 J03-PC1 48 64078 6843 357268 472 0.10679 0.10563 1725 3.3 J03-PC1 49 50116 5850 243906 420 0.11674 0.11653 1904 5.2 J03-PC1 50 25574 2750 132580 192 0.10754 0.10438 1703 3.9 J03-PC1 51 25153 2657 160094 432 0.10562 0.10234 1667 3.9 J03-PC1 52 129173 13773 682940 473 0.10662 0.10662 1742 3.0 J03-PC1 53 56738 5971 301798 478 0.10524 0.10450 1706 2.6 J03-PC1 54 52697 5522 266797 405 0.10478 0.10416 1699 3.3 J03-PC1 55 46506 5404 300538 231 0.11620 0.11603 1896 4.3 J03-PC1 56 21949 2377 355316 292 0.10831 0.10442 1704 5.2 J03-PC1 57 28931 3034 160244 385 0.10487 0.10223 1665 4.2 J03-PC1 58 28761 3055 188177 349 0.10624 0.10362 1690 5.0 J03-PC1 59 41623 4092 199935 190 0.09831 0.09710 1569 4.9 J03-PC1 60 40601 4436 269375 258 0.10925 0.10856 1775 5.4 J03-PC1 61 30758 3192 152466 284 0.10377 0.10139 1650 2.6 J03-PC1 62 45007 4823 226992 265 0.10716 0.10660 1742 3.6 J03-PC1 63 145362 15555 800496 483 0.10701 0.10701 1749 2.2 J03-PC1 64 41036 4197 817704 440 0.10229 0.10125 1647 4.9 J03-PC1 65 50545 5345 566998 215 0.10574 0.10518 1717 5.7 J03-PC1 66 60883 6333 318138 236 0.10402 0.10305 1680 4.0

178 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J03-PC1 67 82293 8763 560408 473 0.10648 0.10352 1688 4.0 J03-PC1 68 65515 7081 1002789 474 0.10808 0.10682 1746 3.6 J03-PC1 69 81484 8878 491355 476 0.10895 0.10599 1731 3.5 J03-PC1 70 105515 11032 619503 473 0.10456 0.09870 1600 3.1 J03-PC1 71 57581 6042 332706 475 0.10492 0.10415 1699 3.7 J03-PC1 72 58499 6145 340136 469 0.10504 0.10422 1701 3.7 J03-PC1 73 30194 3161 191030 473 0.10469 0.10226 1666 3.7 J03-PC1 74 36361 3884 345384 477 0.10683 0.10554 1724 3.3 J03-PC1 75 31471 3270 170085 466 0.10392 0.10165 1655 4.1 J03-PC1 76 64433 6792 336118 472 0.10542 0.10424 1701 3.7 J03-PC1 77 14946 1567 74069 304 0.10482 0.09903 1606 4.0 J03-PC1 78 41703 4363 339079 467 0.10461 0.10375 1692 4.3 J03-PC1 79 40108 4029 245669 291 0.10046 0.09924 1610 3.3 J03-PC1 80 45694 4667 265490 335 0.10213 0.10136 1649 3.0

LC-CC-15: Quartzite from Cebolla Creek, Colorado LL-CC-15 1 63203 7029 337075 477 0.11121 0.10412 1699 3.9 LL-CC-15 2 44915 4947 209479 387 0.11013 0.10296 1678 4.4 LL-CC-15 3 59349 6551 281763 351 0.11038 0.10334 1685 4.6 LL-CC-15 4 108575 24265 293006 332 0.22348 0.21148 2917 4.1 LL-CC-15 5 71209 8179 336190 332 0.11486 0.10753 1758 4.4 LL-CC-15 6 126608 14480 708597 158 0.11437 0.10552 1723 2.7 LL-CC-15 7 208376 48748 586907 233 0.23394 0.21576 2949 4.1 LL-CC-15 8 83619 9379 422407 383 0.11216 0.10467 1708 6.0 LL-CC-15 9 149490 17191 751703 235 0.11500 0.10540 1721 3.3 LL-CC-15 10 68649 7958 334996 190 0.11593 0.10861 1776 4.4 LL-CC-15 11 67894 7733 399772 428 0.11390 0.10666 1743 4.0 LL-CC-15 12 87880 9903 427816 478 0.11269 0.10507 1716 4.1 LL-CC-15 13 102069 11544 524744 475 0.11310 0.10507 1715 3.1 LL-CC-15 14 78600 8890 381974 409 0.11311 0.10570 1726 4.1 LL-CC-15 15 109402 12644 577481 477 0.11557 0.10719 1752 3.0 LL-CC-15 16 136679 16486 1061319 472 0.12062 0.11111 1818 3.7 LL-CC-15 17 78834 8941 373743 476 0.11341 0.10598 1731 3.8 LL-CC-15 18 125629 14300 620697 463 0.11383 0.10504 1715 3.2 LL-CC-15 19 74741 8637 388886 474 0.11555 0.10813 1768 3.8 LL-CC-15 20 110475 13283 684329 309 0.12023 0.11159 1825 2.9 LL-CC-15 21 109192 13693 571025 389 0.12540 0.11655 1904 3.5 LL-CC-15 22 153418 18717 899556 264 0.12200 0.11190 1830 2.8 LL-CC-15 23 117196 13340 629609 363 0.11383 0.10529 1719 4.2

179 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std LL-CC-15 24 41006 4536 225497 313 0.11061 0.10328 1684 4.2 LL-CC-15 25 21129 2306 109107 296 0.10913 0.10124 1647 4.9 LL-CC-15 26 85340 9424 403608 240 0.11042 0.10297 1678 4.2 LL-CC-15 27 40320 4452 197092 333 0.11041 0.10301 1679 4.8 LL-CC-15 28 23171 2684 109225 472 0.11585 0.10772 1761 7.4 LL-CC-15 29 91686 10484 460122 468 0.11435 0.10655 1741 6.5 LL-CC-15 30 44557 5232 210310 202 0.11742 0.11007 1801 7.8 LL-CC-15 31 77530 8706 386978 434 0.11229 0.10494 1713 6.2 LL-CC-15 32 19873 2194 90837 472 0.11041 0.10248 1670 6.3 LL-CC-15 33 45252 8404 142672 332 0.18571 0.17686 2624 5.9 LL-CC-15 34 83601 9643 388179 467 0.11535 0.10772 1761 5.4 LL-CC-15 35 121238 13926 612886 470 0.11486 0.10615 1734 5.1 LL-CC-15 36 79767 9037 395099 470 0.11329 0.10585 1729 5.5 LL-CC-15 37 84462 14819 255563 478 0.17545 0.16549 2513 3.9 LL-CC-15 38 69335 7981 333128 473 0.11510 0.10780 1763 3.8 LL-CC-15 39 66411 7435 341416 475 0.11195 0.10480 1711 3.5 LL-CC-15 40 176578 21161 1033428 472 0.11984 0.10925 1787 3.3 LL-CC-15 41 95763 11242 584301 473 0.11740 0.11087 1814 6.7 LL-CC-15 42 49999 5620 248557 238 0.11241 0.10903 1783 6.4 LL-CC-15 43 36626 4077 182471 476 0.11132 0.10806 1767 5.8 LL-CC-15 44 44548 5172 215603 470 0.11610 0.11271 1844 5.1 LL-CC-15 45 72892 8297 363625 478 0.11383 0.10953 1792 3.8 LL-CC-15 46 139328 15962 765010 478 0.11456 0.10085 1640 3.3 LL-CC-15 47 64733 7907 309969 281 0.12215 0.11815 1928 3.6 LL-CC-15 48 66080 7631 409535 475 0.11548 0.11154 1825 3.0 LL-CC-15 49 84444 9849 497253 470 0.11664 0.11135 1822 3.3 LL-CC-15 50 170557 20165 912312 478 0.11823 0.09662 1560 3.0 LL-CC-15 51 122452 14221 701534 424 0.11614 0.10562 1725 3.1 LL-CC-15 52 100671 11596 593113 472 0.11519 0.10812 1768 2.2 LL-CC-15 53 190106 24646 827017 479 0.12965 0.10039 1631 2.7 LL-CC-15 54 67967 7568 363795 477 0.11135 0.10739 1756 4.4 LL-CC-15 55 75560 8819 422340 473 0.11672 0.11217 1835 3.6 LL-CC-15 56 64788 7220 334091 287 0.11144 0.10763 1760 4.2 LL-CC-15 57 39042 4280 195207 197 0.10963 0.10640 1739 4.9 LL-CC-15 58 149682 17647 731671 476 0.11790 0.10149 1651 3.6 LL-CC-15 59 116254 13497 715445 339 0.11610 0.10667 1743 4.0 LL-CC-15 60 53064 6059 285312 231 0.11418 0.11071 1811 4.0 LL-CC-15 61 96966 12617 419658 474 0.13012 0.12305 2001 3.2 LL-CC-15 62 88213 10026 444486 477 0.11366 0.10809 1767 4.0

180 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std LL-CC-15 63 54427 6238 272315 336 0.11461 0.11111 1818 4.1 LL-CC-15 64 71739 8204 351112 477 0.11436 0.11012 1801 3.7 LL-CC-15 65 76723 8601 397411 476 0.11210 0.10757 1759 4.2 LL-CC-15 66 92113 10451 600138 475 0.11346 0.10748 1757 3.9 LL-CC-15 67 48628 5521 252583 455 0.11353 0.11015 1802 3.5 LL-CC-15 68 80720 9125 390694 478 0.11305 0.10819 1769 3.3 LL-CC-15 69 124290 14134 680800 470 0.11372 0.10304 1680 2.3 LL-CC-15 70 106553 12116 502965 475 0.11371 0.10590 1730 4.6 LL-CC-15 71 102704 12195 506921 285 0.11873 0.11126 1820 4.4 LL-CC-15 72 59215 6958 278854 306 0.11751 0.11381 1861 5.0 LL-CC-15 73 6586 985 32700 192 0.14958 0.14983 2344 9.3 LL-CC-15 74 58340 6330 297296 233 0.10851 0.10501 1715 4.6 LL-CC-15 75 73636 8620 362436 385 0.11705 0.11264 1842 3.5 LL-CC-15 76 178132 21037 931023 474 0.11810 0.09455 1519 3.9 LL-CC-15 77 63255 7075 307659 482 0.11185 0.10811 1768 3.9 LL-CC-15 78 49671 5736 235591 399 0.11548 0.11204 1833 3.7 LL-CC-15 79 128228 14590 643986 475 0.11378 0.10237 1668 3.3 LL-CC-15 80 108878 12573 620101 282 0.11548 0.10726 1753 3.8

J03-NM4: Basal conglomerate of the Uncompaghre Formation, Needle Mountains, Colorado J03-NM4 1 105197 11654 559631 199 0.11078 0.10442 1704 3.8 J03-NM4 2 21198 2306 107584 338 0.10877 0.10471 1709 3.4 J03-NM4 3 12067 1366 63713 307 0.11320 0.10589 1730 8.6 J03-NM4 4 25625 2738 129912 217 0.10683 0.10367 1691 3.6 J03-NM4 5 32630 3724 185959 332 0.11412 0.11265 1843 2.5 J03-NM4 6a 26440 4472 254274 235 0.16916 0.16959 2554 4.6 J03-NM4 6b 13469 1590 98242 122 0.11804 0.11133 1821 4.1 J03-NM4 7 78125 8481 401223 452 0.10856 0.10602 1732 2.7 J03-NM4 8 48512 6699 412747 168 0.13809 0.13847 2208 4.2 J03-NM4 9 82560 8980 427819 387 0.10877 0.10568 1726 2.1 J03-NM4 10 17014 1848 86165 334 0.10863 0.10346 1687 4.8 J03-NM4 11 44942 4753 221878 474 0.10575 0.10512 1716 3.5 J03-NM4 12 21009 2290 108691 169 0.10902 0.10491 1713 3.1 J03-NM4 13 51126 5446 251306 482 0.10652 0.10598 1731 3.0 J03-NM4 14 49136 5363 259610 286 0.10914 0.10870 1778 2.9 J03-NM4 15 28783 3130 140485 304 0.10873 0.10621 1735 3.4 J03-NM4 16 36827 3830 183886 268 0.10401 0.10262 1672 3.7 J03-NM4 17 30939 3452 161436 147 0.11156 0.10957 1792 3.1 J03-NM4 18 59862 6622 313473 239 0.11062 0.10980 1796 3.0

181 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J03-NM4 19 35462 3931 187857 471 0.11086 0.10969 1794 5.1 J03-NM4 20 85811 9312 446517 239 0.10852 0.10502 1715 3.4 J03-NM4 21 28498 3039 141579 212 0.10665 0.10400 1697 4.2 J03-NM4 22 57599 6200 296511 243 0.10764 0.10691 1747 4.0 J03-NM4 23 22125 2290 110165 237 0.10352 0.09963 1617 4.6 J03-NM4 24 45112 5133 266242 335 0.11378 0.11350 1856 4.5 J03-NM4 25 22165 2561 117663 228 0.11555 0.11180 1829 4.1 J03-NM4 26 40713 4382 207341 184 0.10763 0.10685 1746 2.3 J03-NM4 27 17725 1945 87908 399 0.10976 0.10479 1711 5.6 J03-NM4 28 24694 2602 127638 421 0.10539 0.10202 1661 4.0 J03-NM4 29 10963 1199 53963 379 0.10940 0.10156 1653 4.5 J03-NM4 30 55338 6409 352156 206 0.11582 0.11536 1885 3.6 J03-NM4 31 74188 7993 383383 478 0.10775 0.10566 1726 3.3 J03-NM4 32 24619 2525 121209 212 0.10255 0.09912 1608 3.9 J03-NM4 33 23090 2520 117858 472 0.10913 0.10550 1723 4.4 J03-NM4 34 10774 1215 54613 431 0.11278 0.10468 1709 5.3 J03-NM4 35 35238 3765 178529 333 0.10685 0.10539 1721 4.0 J03-NM4 36 99671 10800 512792 475 0.10836 0.10295 1678 2.7 J03-NM4 37 33778 3688 163729 470 0.10919 0.10763 1760 3.3 J03-NM4 38 110980 12191 566419 480 0.10985 0.10282 1676 3.1 J03-NM4 39 74709 8279 395179 480 0.11081 0.10859 1776 3.1 J03-NM4 40 33572 3593 170192 333 0.10703 0.10529 1719 3.8 J03-NM4 41 40448 4324 199442 468 0.10690 0.10607 1733 3.0 J03-NM4 42 26485 2749 135389 194 0.10380 0.10072 1637 4.5 J03-NM4 43 121296 13185 611747 474 0.10870 0.10052 1634 3.8 J03-NM4 44 34535 3659 172198 308 0.10594 0.10431 1702 3.6 J03-NM4 45 82283 9221 432480 474 0.11206 0.10884 1780 3.1 J03-NM4 46 25060 2657 130650 328 0.10604 0.10275 1674 3.2 J03-NM4 47 45243 4952 234786 236 0.10945 0.10899 1783 3.4 J03-NM4 48 96033 11043 660949 424 0.11499 0.10955 1792 2.7 J03-NM4 49 78462 8599 391837 287 0.10959 0.10697 1748 4.0 J03-NM4 50 32465 3504 156043 415 0.10794 0.10602 1732 3.6 J03-NM4 51 28697 3324 154116 115 0.11583 0.11359 1858 2.8 J03-NM4 52 30333 3367 151074 331 0.11100 0.10887 1781 3.8 J03-NM4 53 72865 8023 394741 475 0.11011 0.10812 1768 2.7 J03-NM4 54 31131 3350 155069 474 0.10760 0.10544 1722 3.7 J03-NM4 55 40195 4254 195439 474 0.10583 0.10491 1713 2.8 J03-NM4 56 46065 4965 227343 474 0.10777 0.10726 1753 2.2 J03-NM4 57 40805 4845 235612 261 0.11873 0.11852 1934 3.2

182 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std J03-NM4 58 84039 9049 418622 289 0.10768 0.10445 1705 3.8 J03-NM4 59 22436 2541 114616 313 0.11323 0.10951 1791 3.9 J03-NM4 60 63378 7012 308449 149 0.11063 0.10954 1792 3.3 J03-NM4 61 26279 2777 131355 467 0.10567 0.10260 1672 6.0 J03-NM4 63 20536 2220 104499 472 0.10809 0.10386 1694 6.8 J03-NM4 64 32293 3522 164153 477 0.10907 0.10720 1752 4.2 J03-NM4 65 81053 8812 395696 478 0.10871 0.10581 1728 3.8 J03-NM4 66 40917 4261 204346 390 0.10415 0.10321 1683 3.1 J03-NM4 67 44077 4698 224105 212 0.10658 0.10596 1731 3.5 J03-NM4 68a 25962 2967 131456 194 0.11429 0.11139 1822 4.4 J03-NM4 68b 35184 3744 169981 239 0.10641 0.10492 1713 6.3 J03-NM4 69 24565 2698 123183 370 0.10984 0.10653 1741 5.3 J03-NM4 70 58129 6439 297397 478 0.11077 0.11006 1800 3.5 J03-NM4 71 12926 1466 66123 244 0.11342 0.10656 1741 5.4 J03-NM4 72 79289 8506 398818 432 0.10728 0.10464 1708 3.1 J03-NM4 73 50105 5289 263385 366 0.10555 0.10499 1714 3.4 J03-NM4 74 36268 4137 188313 327 0.11406 0.11319 1851 4.1 J03-NM4 75a 128013 15470 829459 216 0.12084 0.11066 1810 2.8 J03-NM4 75b 143208 15971 789008 266 0.11152 0.10060 1635 3.4 J03-NM4 76 37677 4002 185030 430 0.10621 0.10505 1715 4.3 J03-NM4 77 26080 2769 128147 190 0.10616 0.10306 1680 3.2 J03-NM4 78 7375 808 34978 473 0.10960 0.09848 1596 5.6 J03-NM4 79 18217 1942 89877 308 0.10660 0.10180 1657 3.2

ORT-N: Quartzite of the Ortega Formation, Tusas Mountains, New Mexico ORT-N 1 60440 6573 320610 469 0.10875 0.10219 1664 5.2 ORT-N 2 64165 7284 366577 467 0.11352 0.10655 1741 6.6 ORT-N 3 63578 7142 327747 466 0.11233 0.10545 1722 6.6 ORT-N 4 70333 8104 368094 470 0.11522 0.10790 1764 6.8 ORT-N 5 44972 5056 301873 472 0.11242 0.10649 1740 5.0 ORT-N 6 44421 5040 301651 468 0.11345 0.10751 1758 4.7 ORT-N 7 62355 7700 457583 468 0.12349 0.11609 1897 6.7 ORT-N 8 74574 8089 410666 308 0.10848 0.10136 1649 6.7 ORT-N 9 37155 6853 120065 232 0.18444 0.17601 2616 7.8 ORT-N 10 91699 10380 589771 468 0.11319 0.10518 1718 7.5 ORT-N 11 108047 12459 572279 469 0.11531 0.10659 1742 6.8 ORT-N 12 34249 3975 176816 469 0.11605 0.11083 1813 9.8 ORT-N 13 29900 3311 146377 235 0.11074 0.10628 1737 9.7 ORT-N 14 72811 8109 391052 296 0.11138 0.10417 1700 8.8

183 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std ORT-N 15 117723 13844 639065 269 0.11760 0.10839 1772 7.5 ORT-N 16 31016 3511 160408 331 0.11320 0.10848 1774 6.6 ORT-N 17 40244 4609 224905 463 0.11453 0.10883 1780 9.2 ORT-N 18 65210 7159 340733 257 0.10978 0.10296 1678 8.4 ORT-N 19 67985 7715 356576 470 0.11347 0.10635 1738 7.0 ORT-N 20 124149 13885 652268 469 0.11184 0.10269 1673 8.0 ORT-N 21 48166 5452 250904 469 0.11320 0.10704 1750 5.5 ORT-N 22 43647 5042 211576 472 0.11552 0.10952 1792 6.6 ORT-N 23 29775 3466 148735 472 0.11641 0.11171 1827 5.4 ORT-N 24 24467 2722 124726 426 0.11127 0.10740 1756 8.7 ORT-N 25 59356 6589 307401 478 0.11100 0.10437 1703 5.8 ORT-N 26 67576 7377 366743 477 0.10917 0.10229 1666 5.7 ORT-N 27 40519 4767 202126 433 0.11766 0.11179 1829 8.4 ORT-N 28 67514 7586 355471 471 0.11237 0.10531 1720 4.5 ORT-N 29 98068 11148 553380 475 0.11368 0.10541 1721 4.4 ORT-N 30 37500 4040 189679 478 0.10773 0.10259 1672 6.7 ORT-N 31 96299 11258 587382 421 0.11690 0.10853 1775 4.4 ORT-N 32 76551 8577 400106 475 0.11205 0.10466 1708 4.8 ORT-N 33 70996 7921 377254 467 0.11156 0.10442 1704 4.5 ORT-N 34 97416 11320 489643 383 0.11620 0.10782 1763 3.4 ORT-N 35 31196 3421 159261 477 0.10966 0.10508 1716 4.3 ORT-N 36 40113 4359 203187 360 0.10866 0.10327 1684 5.3 ORT-N 37 29440 3255 192451 293 0.11057 0.10618 1735 4.4 ORT-N 38 82464 9552 420902 471 0.11583 0.10802 1766 4.6 ORT-N 39 109222 13533 608998 470 0.12390 0.11469 1875 3.3 ORT-N 40 67995 7559 348551 467 0.11117 0.10417 1700 4.6 ORT-N 41 51126 5661 259626 285 0.11073 0.10453 1706 3.4 ORT-N 42 53878 5984 268549 462 0.11106 0.10469 1709 4.7 ORT-N 43 54919 6105 278554 474 0.11116 0.10474 1710 3.8 ORT-N 44 65503 7373 358326 289 0.11257 0.10559 1725 4.2 ORT-N 45 85540 9955 463005 336 0.11638 0.10842 1773 4.8 ORT-N 46 114542 13428 553096 473 0.11724 0.10816 1769 2.4 ORT-N 47 63661 7215 332724 470 0.11333 0.10639 1739 4.5 ORT-N 48 89687 10000 455001 477 0.11150 0.10366 1691 3.0 ORT-N 49 72647 8249 443943 474 0.11355 0.10624 1736 4.2 ORT-N 50 59895 6965 289249 470 0.11628 0.10935 1789 4.0 ORT-N 51 84930 9439 454320 474 0.11114 0.10349 1688 3.4 ORT-N 52 117030 13405 599436 475 0.11454 0.10552 1723 2.8 ORT-N 53 76685 8517 398680 429 0.11106 0.10372 1692 3.1

184 RAW COUNTS (cps) 207Pb/206Pb Ratios 207Pb/206Pb Sample # 206Pb 207Pb 238U N raw corrected Age (Ma) Std ORT-N 54 46459 5199 237019 477 0.11190 0.10590 1730 4.2 ORT-N 55 31391 3503 157929 473 0.11161 0.10692 1748 4.1 ORT-N 56 167352 19944 864933 468 0.11918 0.10775 1762 3.3 ORT-N 57 57278 6351 300824 474 0.11087 0.10435 1703 3.3 ORT-N 58 90311 10787 482890 472 0.11944 0.11116 1818 3.6 ORT-N 59 2033 435 13911 200 0.21412 0.28327 3381 7.1 ORT-N 60 111935 12844 597269 441 0.11475 0.10591 1730 2.7 ORT-N 61 57948 6525 344540 433 0.11260 0.10596 1731 3.7 ORT-N 62 91201 10807 467397 466 0.11849 0.11022 1803 3.0 ORT-N 63 167855 19311 836785 468 0.11505 0.10384 1694 2.7 ORT-N 64 54573 6046 270485 401 0.11079 0.10440 1704 3.7 ORT-N 65 394099 49126 1980619 473 0.12465 0.09349 1498 2.4 ORT-N 66 70897 8441 346563 476 0.11906 0.11153 1824 3.6 ORT-N 67 118076 13725 664898 468 0.11624 0.10708 1750 2.9 ORT-N 68 66605 7414 323322 475 0.11131 0.10435 1703 4.3 ORT-N 69 64313 7030 321727 478 0.10931 0.10256 1671 3.3 ORT-N 70 34701 3839 178018 472 0.11064 0.10563 1725 4.2 ORT-N 71 57042 6366 301488 479 0.11161 0.10505 1715 4.9 ORT-N 72 46586 5159 229013 478 0.11074 0.10479 1711 4.5 ORT-N 73 55099 7208 530920 366 0.13081 0.12339 2006 3.6 ORT-N 74 47328 5756 225918 332 0.12162 0.11509 1881 4.7 ORT-N 75 35167 3883 179403 477 0.11043 0.10538 1721 3.8 ORT-N 76 37913 4134 187771 475 0.10905 0.10381 1693 4.7 ORT-N 77 71297 8021 423076 467 0.11250 0.10529 1719 5.0 ORT-N 78 40896 4595 207971 400 0.11235 0.10671 1744 4.4 ORT-N 79 59249 6437 301881 468 0.10864 0.10214 1663 4.2 ORT-N 80 412198 76721 1350734 467 0.18613 0.13880 2212 4.0

# = detrital zircon grain identifier Raw counts (cps) = raw count data presented for each isotope as counts per second (cps) N = number of ratios accepted for each analysis out of 493 total scans 207Pb/206Pb ratios raw = average of accepted 207Pb/206Pb ratios corrected = raw ratio corrected for fractionation and detector non-linearity (see Appendix 2) 207Pb/206Pb Age (Ma) = age of grain calculated using corrected 207Pb/206Pb ratio Std = relative percent error (2σ) on 207Pb/206Pb ratios (see Appendix 2)

185 Bibliography

Afifi, A. M., 1981, Precambrian geology of the Iris area, Gunnison and Saguache counties, Colorado [unpublished M.S. thesis]: Colorado School of Mines, Golden, Colorado, 197 p.

Aleinikoff, J.N., Reed, J.C., Jr., and DeWitt, E., 1993, The Mount Evans batholith in the Colorado Front Range: Revision of its age and reinterpretation of its structure: Geological Society of America Bulletin, v. 105, p. 791 – 806.

Aleinikoff, J.N., Reed, J.C., Jr., and Wooden, J.L., 1995, Lead isotopic evidence for the origin of Paleo- and Mesoproterozoic rocks of the Colorado Province, U.S.A.: Precambrian Research, v. 63, p. 97 – 122.

Anderson, J. L., 1983, Proterozoic anorogenic granite plutonism of North America, in Medaris, L. G., Byers, J., Mickelson, C. W., and Shanks, W. C., eds., Proterozoic Geology: selected papers from an international Proterozoic symposium: Geological Society of America Memoir 161, Boulder, Colorado, p. 133 – 154.

Anderson, J. L., 1987, The origin of A-type Proterozoic magmatism: A model of mantle and crustal overturn: Geological Society of America Abstracts with Programs, v. 19, p. 571.

Anderson, J.L., and Cullers, R.L., 1999, Paleo-and Mesoproterozoic granite plutonism of Colorado and Wyoming: Rocky Mountain Geology, v. 34, pp.149-164.

Anderson, J.L., and Morrison, J., 1992, The role of anorogenic granites in the Proterozoic crustal development of North America, in Condie, K., ed., Proterozoic Crustal Evolution: Amsterdam, The Netherlands, Elsevier, p. 263 – 299.

Barinek, M.F., Foster, C.T., and Chaplinsky, P.P., 1999, Metamorphism and deformation near the approximately 1.4-Ga Mount Ethel Pluton, Park Range, Colorado: Rocky Mountain Geology, v. 34, p. 21-35.

Barker, F., 1969, Precambrian geology of the Needle Mountains, southwestern Colorado: U.S. Geological Survey Professional Paper, p. A1 – A35.

186 Barnes, M.A., Rohs, C.R., Anthony, E.Y., Van Schmus, W.R., and Denison, R.E., 1999, Isotopic and elemental chemistry of subsurface Precambrian igneous rocks, west Texas and eastern New Mexico

Bauer, P.W., and Williams, M.L., 1986, Stratigraphic nomenclature of Proterozoic rocks, northern New Mexico – revisions, redefinitions, and formalization: New Mexico Geology, v. 11, p. 45 – 52.

Bauer, P.W., and Williams, M.L., 1994, The age of Proterozoic orogenesis in New Mexico, U.S.A.: Precambrian Research, v. 67, p. 349 – 356.

Bennett, V.C., and DePaolo, D.J., 1987, Proterozoic crustal history of the western United States as determined by neodymium isotopic mapping: Geological Society of America Bulletin, v. 99, p. 674 – 685.

Bickford, M. E., and Boardman, S. J., 1984, A Proterozoic volcano-plutonic terrane, Gunnison and Salida areas, Colorado: Journal of Geology, v. 92, p. 657 – 666.

Bickford, M.E., and Mose, D.G., 1975, Geochronology of Precambrian rocks in the St. Francois Mountains, southeastern Missouri: Geological Society of America Special Paper 165, p. 48.

Bickford, M. E., Cullers, R. L., Shuster, R. D., Premo, W. R., and Van Schmus, W. R., 1989a, U-Pb zircon geochronology of Proterozoic and Cambrian plutons in the Wet Mountains and southern Front Range, Colorado, in Grambling, J. A., and Tewksbury, B. J., Proterozoic geology of the southern Rocky Mountains: Boulder, Colorado, Geological Society of America Special Paper 235, p. 49 – 64.

Bickford, M. E., Shuster, and Boardman, S. J., 1989b, U-Pb geochronology of the Proterozoic volcano-plutonic terrane in the Gunnison and Salida areas, Colorado, in Grambling, J. A., and Tewksbury, B. J., Proterozoic geology of the southern Rocky Mountains: Boulder, Colorado, Geological Society of America Special Paper 235, p. 33 – 48.

Bowring, S.A., and Karlstrom, K.E., 1990, Growth, stabilization, and reactivation of Proterozoic lithosphere in the southwestern United States: Geology, v. 18, p. 1203 – 1206.

Boyer, R.E., 1962, Petrology and structure of the southern Wet Mountains, Colorado: Geological Society of America Bulletin, v. 73, p. 1047 – 1070.

187 Brasse, H., Lezaeta, P., Rath, V., Schwalenberg, K., Soyer, W., and Haak, V., 2002, The Bolivian Altiplano conductivity anomaly: Journal of Geophysical Research, v. 107,

Brock, M.R., and Singewald, Q.D., 1968, Geologic map of the Mt. Tyndall quadrangle, Custer County, Colorado: U.S. Geological Survey Geologic Quadrangle Map GQ-596.

Bruce, R. M., and Johnson, B. R., 1991, Reconnaissance geologic map of parts of the Zapata Ranch and Mosca Pass quadrangles, Alamosa and Huerfano counties, Colorado: U.S. Geological Survey Miscellaneous Field Studies Map MF-2168, scale 1:24,000.

Burchfiel, B.C., Chen, Z., Hodges, K.V., Liu, Y., Royden, L.H., Deng, C., and Xu, J., 1992, The South Tibetan detachment system, Himalayan Orogen: extension contemporaneous with and parallel to shortening in a collisional mountain belt: Boulder, Colorado, Geological Society of America Special Paper 269, 41 p.

Burchfiel, B.C., Royden, L.H., Tzankov, T., and Nakov, R., 1987, Late Cenozoic extension in Bulgaria; the northern part of the Aegean extensional regime: Geological Society of America Abstracts with Programs, v. 29, p. 377.

Callahan, O., 2002, Preliminary evidence for syntectonic emplacement of Mesoproterozoic plutons in the southern Wet Mountains, Colorado, in Proceedings, 15th Keck Research Symposium in Geology, Amherst, MA, p. 137 – 140.

Cloos, M., 2003, Collisional delamination, lithospheric rupture and magmatism: Geological Society of America Abstracts with Programs, v. 35, p. 95.

Collins, Thomas R. IV, Siddoway, Christine S., Jones, James V. III, and Tellio, Candice, 2004, Shear zone roots in the middle crust: the transition from partitioned deformation to penetrative ductile flow in the northern Wet Mountains, CO: Geological Society of America Abstracts with Programs, v. 36, no. 5.

Collins, W.J., 1994, Upper- and middle-crustal response to delamination: An example from the Lachlan fold belt, eastern Australia: Geology, v. 22, p. 143 – 146.

Condie, K.C., 1982, Plate tectonics model for Proterozoic continental accretion in the southwestern United States: Geology, v. 10, p. 37 – 42.

188 Condie, K.C., 1986, Geochemistry and tectonic setting of early Proterozoic supracrustal rocks in the southwestern United States: Journal of Geology, v. 94, p. 845 – 864.

Condie, K.C., and Nuter, J.A., 1981, Geochemistry of the Dubois greenstone succession: an early Proterozoic bimodal volcanic association in west- central Colorado: Precambrian Research, v. 15, p. 131 – 155.

Cox, R., Gutmann, Ethan D., and Hines, P.G., 2002a, Diagenetic origin for quartz-pebble conglomerate: Geology, v. 30, p. 323 – 326.

Cox, R., Martin, M.W., Comstock, J.C., Dickerson, L.S., Ekstrom, I.L., and Sammons, J.H., 2002b, Sedimentology, stratigraphy, and geochronology of the Proterozoic Mazatzal Group, central Arizona: Geological Society of America Bulletin, v. 114, p. 1535 – 1549.

Cullers, R.L., Griffin, T., Bickford, M.E., and Anderson, J.L., 1992, Origin and chemical evolution of the 1360 Ma San Isabel Batholith, Wet Mountains, Colorado: a mid-crustal granite of anorogenic affinities: Geological Society of America Bulletin, v. 104, p. 316 – 328.

Cullers, R.L., Koch, R.J., and Bickford, M.E., 1981, Chemical evolution of magmas in the Proterozoic terrane of the St. Francois Mountains, southeastern Missouri, 2: Trace element data: Journal of Geophysical Research, v. 86, p. 10388 – 10401.

Cullers, R.L., Stone, J., Anderson, J.L., Sassarini, N., and Bickford, M.E., 1993, Petrogenesis of Mesoproterozoic Oak Creek and West McCoy Gulch plutons, Colorado: an example of cumulate unmixing of a mid-crustal, two-mica granite of anorogenic affinity: Precambrian Research, v. 62, p. 139 – 169.

Davis, D. W., 1982, Optimum linear regression and error estimation applied to U- Pb data: Canadian Journal of Earth Sciences, v. 19, p. 2141 – 2149.

Doe, M.F., and Karlstrom, K.E., 1991, Structural geology of an early Proterozoic foreland thrust belt, Mazatzal Mountains, Arizona, in Karlstrom, K.E., ed., Proterozoic geology and ore deposits of Arizona: Flagstaff, Arizona, Arizona Geological Society Digest, v. 19, p. 181 – 192.

189 Dott, R.H., Jr., 1983, The Proterozoic red quartzite enigma in the north-central United States: resolved by plate collision? in Medaris, L.G., ed., Early Proterozoic geology of the Great Lakes region: Boulder, Colorado, Geological Society of America Memoir, v. 160, p. 129 – 141.

Dott, R.H., Jr., 2003, The importance of eolian abrasion in supermature quartz sandstones and the paradox of weathering on vegetation-free landscapes: Journal of Geology, v. 111, p. 387 – 405.

Dott, R.H., Jr., and Dalziel, I.W.D., 1972, Age and correlation of the Precambrian Baraboo Quartzite of Wisconsin: Journal of Geology, v. 80, p. 552 – 568.

Duebendorfer, E. M., Chamberlain, K. R., and Jones, C. S., 2001, Paleoproterozoic tectonic history of the Cerbat Mountains, northwestern Arizona: implications for crustal assembly in the southwestern United States: Geological Society of America Bulletin, v. 113, p. 575 – 590.

Emslie, R.F., 1978, Anorthosite massifs, rapakivi granites, and late Proterozoic rifting of North America: Precambrian Research, v. 7, p. 61 – 98.

Ferguson, C.B., Duebendorfer, E.M., and Chamberlain, K.R., 2004, Synkinematic intrusion of the 1.4-Ga Boriana Canyon Pluton, Northwestern Arizona: Implications for ca. 1.4-Ga regional strain in the Western United States: Journal of Geology, v. 112, p. 165 – 183.

Frost, B. R., Chamberlain, K. R., and Schumacher, J. C., 2000, Sphene (titanite): phase relations and role as a geochronometer: Chemical Geology, v. 172, p. 131 – 148.

Frost, C. D., and Frost, B. R., 1997, Reduced rapakivi-type granites: The tholeiite connection: Geology, v. 25, p. 647 – 650.

Frost, C.D., Bell, J.M., Frost, B.R., and Chamberlain, K.R., 2001, Crustal growth by magmatic underplating: Isotopic evidence from the northern Sherman batholith: Geology, v. 29, p. 515 – 518.

Gibson, R.G., and Simpson, C., 1988, Proterozoic polydeformation in basement rocks of the Needle Mountains, Colorado: Geological Society of America Bulletin, v. 100, p. 1957 – 1970.

Giles, D., and Betts, P., 2000, The early to middle Proterozoic configuration of Australia and its implications for Australian-US relations: Abstracts – Geological Society of Australia, v. 59, p. 174.

190 Givot, R.M., 1998, Proterozoic deformation and metamorphism of Five Points Gulch, Colorado [Bachelor’s thesis]: Colorado Springs, Colorado College, 79 p.

Givot, R.M., and Siddoway, C.S., 1998, Deformation of mid-Proterozoic metamorphic rocks, Five Points Gulch shear zone, Colorado: Geological Society of America Abstracts with Programs, v. 30, no. 6, p. 9.

Graubard, C.M., and Mattinson, J.M., 1990 Syntectonic emplacement of the approximately 1440 Ma Mt. Evans Pluton and history of motion along the Idaho Springs-Ralston Creek shear zone, central Front Range, Colorado: Geological Society of America Abstracts with Programs, v. 22, p. 12.

Harris, C.W., 1990, Polyphase suprastructure deformation in metasedimentary rocks of the Uncompaghre Group; remnant of an early Proterozoic fold belt in Southwest Colorado: Geological Society of America Bulletin, v. 102, p. 664 – 678.

Harris, C.W., Gibson, R.G., Simpson, C., and Eriksson, K.A., 1987, Proterozoic cuspate basement-cover structure, Needle Mountains, Colorado: Geology, v. 15, p. 950 – 953.

Henry, D. M., 1975, Sedimentology and stratigraphy of the Baraboo Quartzite of south-central Wisconsin [M.S. thesis]: Madison, Wisconsin, University of Wisconsin, 90 p.

Hill, B.M., and Bickford, M.E., 2001, Paleoproterozoic rocks of central Colorado: Accreted arcs or extended older crust?: Geology, v. 29, p. 1015 – 1018.

Hoffman, P.F., 1988, United Plates of America, the birth of a craton: Early Proterozoic assembly and growth of Laurentia: Annual Reviews of Earth and Planetary Sciences, v. 16, p. 543 – 603.

Hoffman, P. F., 1989, Speculations on Laurentia’s first gigayear (2.0 – 1.0 Ga): Geology, v. 17, p. 135 – 138.

Holcombe, R. J., 2003, GEOrient v. 9.1, http://www.holcombe.net.au/software/ index.html

Holm, D., Schneider, D., and Coath, C.D., 1998, Age and deformation of Early Proterozoic quartzites in the southern Lake Superior region: Implications for extent of foreland deformation during final assembly of Laurentia: Geology, v. 26, p. 907 – 910.

191 Houseman, G.A., McKenzie, D.P., and Molnar, P., 1981, Convective instability of a thickened boundary layer and its relevance for the thermal evolution of continental convergent belts: Journal of Geophysical Research, v. 86, p. 6115 – 6132.

Jaffey, A.H., Flynn, K.F., Glendenin, L.E., Bentley, W.C., and Essling, A.M., 1971, Precision measurements of halflives and specific activities of 238U and 235U: Physical Reviews, v. 4, p. 1889 – 1906.

Jessup, M.J., 2003, Tectonic history of Proterozoic rocks in the Black Canyon of the Gunnison, Colorado [M.S. thesis]: Albuquerque, New Mexico, University of New Mexico, 143 p.

Jessup, M.J., Karlstrom, K.E., Livaccari, R., Connelly, J., Tyson, A., and Rogers, S.A., 2005, Complex Proterozoic crustal assembly of southwestern North America in an arcuate subduction system: The Black Canyon of the Gunnison, southwestern Colorado: in Karlstrom, K.E., and Keller, G.R., eds., The Rocky Mountain Region – An Evolving Lithosphere: Tectonics, Geochemistry, and Geophysics: American Geophysical Union Monograph.

Johnson, R. B., 1969, Geologic map of the Trinidad quadrangle, south-central Colorado: U.S. Geological Survey Miscellaneous Investigations Series Map I-558, scale 1:250,000.

Johnson, B. R., and Bruce, R. M., 1991, Reconnaissance geologic map of parts of the Twin Peaks and Blanca Peak quadrangles, Alamosa, Costilla, and Huerfano counties, Colorado: U.S. Geological Survey Miscellaneous Field Studies Map MF-2169, scale 1:24,000.

Johnson, B. R., Lindsey, D. A., Bruce, R. M., and Soulliere, S. J., 1987, Geologic map of the Sangre de Cristo Wilderness Study Area, Colorado: U.S. Geological Survey Miscellaneous Field Studies Map MF-1635-B, scale 1:62,500.

Johnsson, M.J., Stallard, R.F., and Lundberg, N., 1991, Controls on the composition of fluvial sands from a tropical weathering environment; sands of the Orinoco River drainage basin, Venezuela and Colombia: Geological Society of America Bulletin, v. 103, p. 1622 – 1647.

Johnsson, M.J., Stallard, R.F., and Meade, R.H., 1988, First-cycle quartz arenites in the Orinoco River basin, Venezuela and Colombia: Journal of Geology, v. 96, p. 263 – 277.

192 Karlstrom, K. E., 1989, Early recumbent folding during Proterozoic orogeny in central Arizona, in Grambling, J. A. and Tewksbury, B. A., eds., Proterozoic geology of the southern Rocky Mountains: Boulder, Colorado, Geological Society of America Special Paper, v. 235, p. 155 – 171.

Karlstrom, K. E., and Bowring, S. A., 1988, Early Proterozoic assembly of tectonostratigraphic terranes in southwestern North America: Journal of Geology, v. 96, p. 561 – 576.

Karlstrom, K. E., and Bowring, S. A., 1993, Proterozoic orogenic history in Arizona, in Reed, J.C., Jr., Bickford, M.E., Houston, R.S., Link, P.K., Rankin, D.W., Sims, P.K., and Van Schmus, W.R., eds., Precambrian: Conterminous U.S.: Boulder, Colorado, Geological Society of America, The Geology of North America, v. C-2, p. 188 – 211.

Karlstrom, K.E., and Daniel, C.G., 1993, Restoration of Laramide right-lateral strike slip in northern New Mexico by using Proterozoic piercing points: tectonic implications from the Proterozoic to the Cenozoic: Geology, v. 21, p. 1139 – 1142.

Karlstrom, K.E., and Humphreys, E.D., 1998, Persistent influence of Proterozoic accretionary boundaries in the tectonic evolution of southwestern North America; interaction of cratonic grain and mantle modification events: Rocky Mountain Geology, v. 33, p. 161 – 179.

Karlstrom, K.E., Ahall, K.I., Harlan, S.S., Williams, M.L., McLelland, J., and Geissman, J.W., 2001, Long-lived (1.8-1.0 Ga) convergent orogen in southern Laurentia, its extensions to Australia and Baltica, and implications for refining Rodinia: Precambrian Research, v. 111, p. 5 – 30.

Karlstrom, K.E., Amato, J.M., Williams, M.L., Heizler, M., Shaw, C., Read, A., and Bauer, P., 2004, Proterozoic tectonic evolution of the New Mexico region: a synthesis, in Mack, G.H., and Giles, K.A., eds., The Geology of New Mexico: A Geologic History: Albuquerque, New Mexico, New Mexico Geological Society Special Publication 11, p. 1 – 34.

Kelley, S.A., and Chapin, C.E., 2004, Denudation history and internal structure of the Front Range and Wet Mountains, Colorado, based on apatite-fission- track thermochronology: New Mexico Bureau of Geology & Mineral Resources Bulletin 160, p. 41 – 77.

193 Kirby, E., Karlstrom, K.E., Andronicos, C.L., and Dallmeyer, R.D., 1995, Tectonic setting of the Sandia Pluton; an orogenic 1.4 Ga granite in New Mexico: Tectonics, v. 14, p. 185 – 201.

Kluth, C. F., 1986, Plate tectonics of the Ancestral Rocky Mountains, in Peterson, J. A., ed., Paleotectonics and sedimentation in the Rocky Mountain region, United States: American Association of Petroleum Geologists Memoir 41, p. 353 – 369.

Kluth, C. F., and Coney, P. J., 1981, Plate tectonics of the Ancestral Rocky Mountains: Geology, v. 9, p. 10 – 15.

Knoper, M.W., and Condie, K.C., 1988, Geochemistry and petrogenesis of early Proterozoic amphibolites, west-central Colorado, U.S.A.: Chemical Geology, v. 67, p. 209 – 225.

Krogh, T.E., 1973, A low-contamination method for hydrothermal decomposition of zircon and extraction of U and Pb for isotopic age determination: Geochimica et Cosmochimica Acta, v. 37, p. 485 – 494.

Krogh, T. E., 1982, Improved accuracy of U-Pb zircon ages by the creation of more concordant systems using an abrasion technique: Geochimica et Cosmochimica Acta, v. 46, p. 637 – 649.

Lanzirotti, A., 1988, Geology and geochemistry of a Proterozoic supracrustal and intrusive sequence in the Central Wet Mountains, Colorado [M.S. thesis]: Socorro, New Mexico, New Mexico Institute of Mining and Technology, 164 p.

Lindsey, D. A., Johnson, B. R., Soulierre, S. J., Bruce, R. M., and Hafner, K., 1986, Geologic map of the Beck Mountain, Crestone Peak, and Crestone quadrangles, Custer, Huerfano, and Saguache counties, Colorado: U.S. Geological Survey Miscellaneous Field Studies Map MF-1878, scale 1:24,000.

Lindsey, D. A., Soullierre, S. J., Hafner, K., and Flores, R. J., 1985, Geologic map of Rito Alto Peak and northeastern part of Mirage quadrangles, Custer and Saguache counties, Colorado: U.S. Geological Survey Miscellaneous Studies Map MF-1787, scale 1:24,000.

Loiselle, M.C., and Wones, D.R., 1979, Characteristics and origin of anorogenic granites: Geological Society of America Abstracts with Programs, v. 11, no. 7, p. 468.

194 McCoy, A.M., Karlstrom, K.E., Williams, M.L., and Shaw, C.A., Proterozoic ancestry of the Colorado mineral belt: ca. 1.4 Ga shear zone system in southern Colorado, in Karlstrom, K.E., and Keller, G.R., eds., Lithospheric Structure and Evolution of the Rocky Mountain Region: Washington, D.C., American Geophysical Union Monograph xx, p.

Medaris, L.G., Jr., Singer, B.S., Dott, R.H., Naymark, A., Johnson, C.M., and Schott, R.C., 2003, Late Paleoproterozoic climate, tectonics, and metamorphism in the Southern Lake Superior Region and Proto-North America: Evidence from Baraboo Interval Quartzites: Journal of Geology, v. 111, p. 243 – 257.

Menuge, J.F., Brewer, T.S., and Seeger, C.M., 2002, Petrogenesis of metaluminous A-type rhyolites from the St. Francois Mountains, Missouri, and the Mesoproterozoic evolution of the southern Laurentian margin: Precambrian Research, v. 113, p. 269 – 291.

Nelson, K.D., and Project INDEPTH, 1996, Partially molten middle crust beneath southern Tibet: Synthesis of Project INDEPTH: Science, v. 274, p. 1684 – 1688.

Nyman, M.W., and Karlstrom, K.E., 1997, Pluton emplacement processes and tectonic setting of the 1.42 Ga Signal Batholith, SW USA; important role of crustal anisotropy during regional shortening: Precambrian Research, v. 82, p. 237 – 263.

Nyman, M.W., Karlstrom, K.E., Kirby, E., and Graubard, C.M., 1994, Mesoproterozoic contractional orogeny in western North America; evidence from ca. 1.4 Ga plutons: Geology, v. 22, p. 901 – 904.

Patchett, P.J., and Ruiz, J., 1989, Nd isotopes and the origin of Grenville-age rocks in Texas: Implications for Proterozoic evolution of the United States mid-continent region: Journal of Geology, v. 97, p. 685 – 695.

Perkins, G., 2002, Dynamic high-temperature metamorphism and deformation during emplacement of the San Isabel Pluton, Wet Mountains, Colorado, in Proceedings, 15th Keck Research Symposium in Geology, Amherst, MA, p. 153 – 156.

Pidgeon, R. T., Bosch, D., and Bruguier, O., 1996, Inherited zircon and titanite U- Pb systems in an Archaean syenite from southwestern Australia: implications for U-Pb stability of titanite: Earth and Planetary Science Letters, v. 141, 187 – 198.

195 Puls, D. D., 1986, Geometric and kinematic analysis of a Proterozoic foreland thrust belt, northern Mazatzal Mountains, central Arizona [unpublished M. S. thesis]: Flagstaff, Arizona, Northern Arizona University, 102 p.

Raymond, W.H., Leiggi, P.A., and Sheridan, D.M., 1980, Sapphirine in Precambrian rocks associated with stratabound sulfide deposits, Custer County, Colorado: Washington, D.C., U.S. Geological Survey Bulletin 1513, 22 p.

Reed, J. C., Jr., Bickford, M. E., and Tweto, O., 1993, Proterozoic accretionary terranes of Colorado and southern Wyoming, in Reed, J. C., Jr., ed., Precambrian: Conterminous U.S., Volume C-2: The Geology of North America: Boulder, Colorado, Geological Society of America, p. 110 – 121.

Reuss, R.L., 1970, Geology and petrology of the Wilson Park area, Fremont County, Colorado [PhD dissertation]: Ann Arbor, Michigan, University of Michigan, 184 p.

Reuss, R.L., 1974, Precambrian quartzite-schist sequence in Wilson Park, Fremont County, Colorado: Mountain Geologist, v. 11, p. 45 – 58.

Rivers, T. 1997, Lithotectonic elements of the Grenville Province: review and tectonic implications: Precambrian Research, v. 86, p. 117 – 154.

Rogers, S.A., Connelly, J.N., Jones, J.V. III, and Rhea, C., 2004, U-Pb zircon geochronology of Proterozoic granites of the Taylor River area, SW Sawatch Range, Gunnison County, central Colorado: Geological Society of America Abstracts with Programs, v. 36, no. 5.

Romano, D., Holm, D.K., and Foland, K.A., 2000, Determining the extent and nature of Mazatzal-related overprinting of the Penokean orogenic belt in the southern Lake Superior region, north-central USA: Precambrian Research, v. 104, p. 25 – 46.

Royden, L., 1996, Coupling and decoupling of crust and mantle in convergent orogens; implications for strain partitioning in the crust: Journal of Geophysical Research, v. 101, p. 17679 – 17705.

Sabin, T. G., 1994, Geochronology and isotope geochemistry of early Proterozoic rocks exposed in portions of the Twin Peaks and Blanca Peak quadrangles, Alamosa, Costilla, and Huerfano counties, Colorado [unpublished M. S. thesis]: Lawrence, Kansas, University of Kansas, 94 p.

196 Scott, G. R., Taylor, R. B., Epis, R. C., and Wobus, R. A., 1978, Geologic map of the Pueblo 1° x 2° quadrangle, south-central Colorado: U.S. Geological Survey Miscellaneous Investigations Series Map I-1022, scale 1:250,000.

Sears, J.W., and Price, R.A., 2000, New look at the Siberian connection: no SWEAT: Geology, v. 28, p. 423 – 426.

Selverstone, J., Hodgins, M., Aleinikoff, J.N., and Fanning, C.M., 2000, Mesoproterozoic reactivation of a Paleoproterozoic transcurrent boundary in the northern Colorado Front Range: Implications for ~1.7- and 1.4-Ga tectonism: Rocky Mountain Geology, v. 35, p. 139 – 162.

Siddoway, C.S., Givot, R.M., Bodle, C.D., and Heizler, M.T., 2000, Dynamic versus anorogenic setting for Mesoproterozoic plutonism in the Wet Mountains, Colorado: Does the interpretation depend on level of exposure? Rocky Mountain Geology, v. 35, p. 91 – 111.

Siddoway, C.S., Jones, J., Williams, M.L., Andronicos, C.L., Connelly, J., and Karlstrom, K.E., 2002, The Mesoproterozoic record in the Wet Mountains, Colorado: Geological Society of America Abstracts with Programs, v. 34, p. 181.

Silver, L.T., 1965, Mazatzal orogeny and tectonic episodicity: Geological Society of America Special Paper, p. 185 – 188.

Silver, L.T., and Barker, F., 1968, Geochronology of Precambrian rocks of the Needle Mountains, southwestern Colorado: Part 1, U-Pb zircon results: Special Paper – Geological Society of America, p. 204 – 205.

Silver, L.T., Conway, C.M., and Ludwig, K.R., 1986, Implications of a precise chronology for Early Proterozoic crustal evolution and caldera formation in the Tonto basin-Mazatzal Mountains region, Arizona: Geological Society of America Abstract with Programs, v. 18, p. 413.

Shaw, C.A., and Karlstrom, K.E., 1999, The Yavapai-Mazatzal crustal boundary in the Southern Rocky Mountains: Rocky Mountain Geology, v. 34, p. 37 – 52.

Shaw, C.A., Heizler, M.T., and Karlstrom, K.E., 2005, 40Ar/30Ar thermochronologic record of 1.45-1.35 Ga intracontinental tectonism in the southern Rocky Mountains: Interplay of conductive and advective heating with intracontinental deformation, in AGU monograph

197 Shaw, C.A., Karlstrom, K.E., Williams, M.L., Jercinovic, M.J., and McCoy, A.M., 2001, Electron-microprobe monazite dating of ca. 1.71 – 1.63 Ga and ca. 1.45 – 1.38 Ga deformation in the Homestake shear zone, Colorado: Origin and early evolution of a persistent intracontinental tectonic zone: Geology, v. 29, p. 739 – 742.

Shaw, C.A., Snee, L.W., Selverstone, J., and Reed, J.C., Jr., 1999, 40Ar/39Ar thermochronology of Mesoproterozoic metamorphism in the Colorado Front Range: Journal of Geology, v. 107, p. 49 – 67.

Soegaard, K., and Eriksson, K.A., 1985, Evidence of tide, storm, and wave interaction on a Precambrian siliciclastic shelf: The 1,700 M.Y. Ortega Group, New Mexico: Journal of Sedimentary Petrology, v. 55, p. 672 – 684.

Soegaard, K., and Eriksson, K.A., 1989, Origin of thick, first-cycle quartz arenite successions: evidence from the 1.7 Ga Ortega Group, northern New Mexico: Precambrian Research, v. 43, p. 129 – 141.

Stacey, J. S., and Kramers, J. D., 1975, Approximation of terrestrial lead isotopic evolution by a two-stage model: Earth and Planetary Science Letters, v. 26, p. 207 – 221.

Taylor, R. B., Scott, G. R., and Wobus, R. A., 1975, Reconnaissance geologic map of the Howard quadrangle, central Colorado: U. S. Geological Survey Miscellaneous Investigations Map I-892, scale 1:62,500.

Taylor, R.B., Scott, G.R., Wobus, R.A., and Epis, R.C., 1975a, Reconnaissance geologic map of the Cotopaxi 15-minute Quadrangle, Fremont and Custer Counties, Colorado: U.S. Geological Survey Miscellaneous Investigations Series Map I-900, scale 1:62,500.

Taylor, R.B., Scott, G.R., Wobus, R.A., and Epis, R.C., 1975b, Reconnaissance geologic map of the Royal Gorge Quadrangle, Fremont and Custer Counties, Colorado: U.S. Geological Survey Miscellaneous Investigation Series Map I-869, scale 1:62,500.

Trevena, A.S., 1979, Studies in sandstone petrology: Origin of the Precambrian Mazatzal Quartzite and provenance of detrital feldspar [Ph.D. thesis]: Salt Lake City, Utah, University of Utah, 390 p.

Tweto, O., 1987, Rock units of the Precambrian basement in Colorado: U.S. Geological Survey Professional Paper 1321-A, 54 p.

198 Tweto, O., 1979, Geologic map of Colorado: Colorado Geological Survey, scale 1:500,000.

Tweto, O., and Sims, P.K., 1963, Precambrian ancestry of the Colorado mineral belt: Geological Society of America Bulletin, v. 74, p. 991 – 1014.

Van Schmus, W.R., MacNeill, L.C., Holm, D.K., and Boerboom, T.J., 2001, New U-Pb ages from Minnesota, Michigan and Wisconsin; implications for late Paleoproterozoic crustal stabilization: Proceedings and Abstracts – Institute on Lake Superior Geology, v. 47, p. 100 – 101.

Van Schmus, W.R., Thurman, M.E., and Peterman, Z.E., 1975, Geology and Rb- Sr chronology of Middle Precambrian rocks in eastern and central Wisconsin: Geological Society of America Bulletin, v. 86, p. 1255 – 1265.

Van Schmus, W.R., Bickford, M.E., Turek, A., 1996, Proterozoic geology of the east-central Midcontinent basement, in van der Pluijm, B.A., and Catacosinos, P., eds., Basement and Basins of Eastern North America: Boulder, Colorado, Geological Society of America Special Paper 308, p. 7 – 32.

Van Wyck, N., 1995, Major and trace element, common Pb, Sm-Nd, and zircon geochronology constraints on petrogenesis and tectonic setting of pre- and Early Proterozoic rocks in Wisconsin [Ph.D. thesis]: Madison, Wisconsin, University of Wisconsin, 280 p.

Verts, L. A., Chamberlain, K. R., and Frost, C. D., 1996, U-Pb sphene dating of metamorphism: the importance of sphene growth in the contact aureole of the Red Mountain pluton, Laramie Mountains, Wyoming: Contributions to Mineralogy and Petrology, v. 125, p. 186 – 199.

Wells, J. D., Sheridan, D. M., and Albee, A. L., 1964, Relationship of Precambrian quartzite-schist sequence along Coal Creek to Idaho Springs Formation, Front Range, Colorado: U.S. Geological Survey Professional Paper 454-O, 22 p.

Williams, M.L., 1991, Heterogeneous deformation in a ductile fold-thrust belt: The Proterozoic structural history of the Tusas Mountains, New Mexico: Geological Society of America Bulletin, v. 102, p. 171 – 188.

199 Williams, M.L., and Karlstrom, K.E., 1996, Looping P-T paths and high-T, low-P middle crustal metamorphism: Proterozoic evolution of the southwestern United States: Geology, v. 24, p. 1119 – 1122.

Williams, M.L., Karlstrom, K.E., Jessup, M., Jones, J., and Connelly, J., 2003, Proterozoic rhyolite-quartzite sequences of the Southwest: Syntectonic "cover" and stratigraphic breaks (~1695 and ~1660 Ma) between orogenic pulses: Geological Society of America Abstracts with Programs, v. 35, no. 5,

Windley, B., 1992, Proterozoic collisional and accretionary orogens, in Condie, K.C., ed., Proterozoic crustal evolution: Amsterdam, Netherlands, Elsevier, p. 419 – 446.

Wooden, J.L., Stacey, J.S., Howard, K.A., Doe, B.R., and Miller, D.M., 1988, Pb isotopic evidence for the formation of Proterozoic crust in the southwestern United States, in Ernst, W.G., ed., Metamorphism and crustal evolution of the western United States [Rubey Volume]: Englewood Cliffs, New Jersey, Prentice-Hall, v. 7, p. 68 – 86.

Wooden, J.L., and DeWitt, E., 1991, Pb isotopic evidence for the boundary between the early Proterozoic Mojave and central Arizona crustal provinces in western Arizona, in Karlstrom, K.E., ed., Proterozoic geology and ore deposits of Arizona: Tuscon, Arizona, Arizona Geological Society Digest, v. 19, p. 27 – 50.

200 Vita

James V. Jones III was born September 30, 1975, in New Orleans, Louisiana, to James V. Jones, Jr., and Bridget Posey Jones. He grew up in Ruston, Louisiana, and graduated from Cedar Creek School in May 1993. He attended The University of the South in Sewanee, Tennessee, and graduated with a B.S. cum laude in Geology in May 1997. Upon graduation, he enrolled at University of Wyoming and earned an M.S. in Geology in December 1999. His master’s research was supervised by Dr. Arthur W. Snoke and focused on the Mesozoic tectonic history of the Ruby Mountains metamorphic core complex in northeastern Nevada. He enrolled at The University of Texas at Austin in August 2000 and has conducted his PhD research on the Precambrian tectonic history of southern Colorado under the supervision of Dr. James N. Connelly. This research has led to the preparation of three manuscripts to be submitted to peer-reviewed journals in 2005, and he is also a coauthor on one additional manuscript. During his time at U.T. Austin, the author has been a graduate teaching assistant in the undergraduate mineralogy and petrology curriculum and in summer field course and a graduate research assistant in Jim Connelly’s U-Pb geochronology lab. He has also served as part-time faculty at Trinity University in San Antonio, Texas, teaching Earth Materials.

Permanent address: 4109 Avenue A Austin, TX 78751 This dissertation was typed by the author.

201 Back Plates

The following two oversized color plates are formatted to print on tabloid- sized (11”x17”) paper and may require a special printer or plotter. The captions for both plates are included below and on the plates themselves.

Plate 1. Geologic map of the Music Pass area. See Figure 1.3 for explanation of map symbols. Contacts modified from Lindsey et al. (1986). Lower-hemisphere, equal-area stereonet diagrams represent poles to foliation in host-rock gneisses (A) and in the Music Pass quartz monzonite (C) and lineation data from gneissic rocks (B). All structural data and analysis are from this study. Interpreted fabric elements were determined primarily based on orientation and are discussed in the text along with evidence for relative timing relationships. Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003).

Plate 2. Generalized geologic map of the eastern Arkansas River Gorge (ARG). See index for sources of geologic mapping and structural data. New U-Pb ages (this study) and published ages (Crampton Mt. batholith; Bickford et al., 1989) are indicated. Structural data and analysis from the Five Points Gulch shear zone (FPSZ) and Parkdale domain (this study) are represented on lower-hemisphere, equal-area stereonet diagrams. See text for details. Foliation (plotted as poles) and lineation data are plotted separately, and the rock type from which measurements were taken is indicated to the left of the diagrams. Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003).

202 Plate 1. Geologic map of the Music Pass area. See Figure 1.3 for explanation of map symbols. Contacts modified from Lindsey et al. (1986). Lower-hemisphere, equal-area stereonet diagrams represent poles to foliation in host-rock gneisses (A) and in the Music Pass quartz monzonite (C) and lineation data from gneissic rocks (B). All structural data and analysis are from this study. Interpreted fabric elements were determined primarily based on orientation and are discussed in the text along with evidence for relative timing relationships. Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003).

203 Plate 2. Generalized geologic map of the eastern Arkansas River Gorge (ARG). See index for sources of geologic mapping and structural data. New U-Pb ages (this study) and published ages (Crampton Mt. batholith; Bickford et al., 1989) are indicated. Structural data and analysis from the Five Points Gulch shear zone (FPSZ) and Parkdale domain (this study) are represented on lower-hemisphere, equal-area stereonet diagrams. See text for details. Foliation (plotted as poles) and lineation data are plotted separately, and the rock type from which measurements were taken is indicated to the left of the diagrams. Average orientations were calculated using GEOrient 9.1 (Holcombe, 2003).

204