Tectonic Evolution of the Northern Sierra Nevada

Home , Batholith, Blue Wing Mountains, Cathedral Range, Jackson Mountains, Sahwave Mountains, Seven Troughs Range

TECTONIC EVOLUTION OF THE

NORTHERN SIERRA NEVADA BATHOLITH

A DISSERTATION

SUBMITTED TO THE DEPARTMENT OF

GEOLOGICAL AND ENVIRONMENTAL SCIENCES

AND THE COMMITTEE ON GRADUATE STUDIES

OF STANFORD UNIVERSITY

IN PARTIAL FULFILLMENT OF THE REQUIREMENTS

FOR THE DEGREE OF

DOCTOR OF PHILOSOPHY

Nicholas James Van Buer

December 2011

This work is licensed under a Creative Commons Attribution- Noncommercial 3.0 United States License. http://creativecommons.org/licenses/by-nc/3.0/us/

This dissertation is online at: http://purl.stanford.edu/xb187vq0064

Includes supplemental files: 1. Plate 1. Geologic Map of the Jayhawk Well 7.5' Quadrangle, Pershing County, Nevada (jayhawkwell.pdf) 2. Plate 2. Geologic Map of the Juniper Pass 7.5' Quadrangle, Pershing County, Nevada (Juniperpass.pdf) 3. Plate 3. Geologic Map of the Tohakum Peak NE 7.5' Quadrangle, Pershing County, Nevada (TohakumpkNE.pdf) 4. Plate 4. Geologic Map of the Tunnel Spring 7.5' Quadrangle, Pershing County, Nevada (tunnelspr.pdf) 5. Plate 5. Geologic Map of the Bob Spring 7.5' Quadrangle, Pershing County, Nevada (bobspring.pdf) 6. Plate 6. Geologic Map of the Tohakum Peak SE 7.5' Quadrangle, Pershing County, Nevada (TohakumpkSE.pdf) 7. Plate 7. Geologic Map of the Sage Hen Spring 7.5' Quadrangle, Pershing County, Nevada (SageHenSpr.pdf) 8. Plate 8. Geologic Map of the Bluewing Spring 7.5' Quadrangle, Pershing County, Nevada (BluewingSpr.pdf)

ii I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

Elizabeth Miller, Primary Adviser

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

Trevor Dumitru

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

Martin Grove

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

Gail Mahood

Approved for the Stanford University Committee on Graduate Studies. Patricia J. Gumport, Vice Provost Graduate Education

This signature page was generated electronically upon submission of this dissertation in electronic format. An original signed hard copy of the signature page is on file in University Archives.

iii iv

ABSTRACT

About a third of the Mesozoic Sierra Nevada Batholith actually extends northward out of the Sierra Nevada proper into the Basin and Range Province of northwest Nevada, where it is obscured by late Cenozoic volcanic rocks, extensional faulting, and sediment-filled grabens. This segment differs from the main Sierra Nevada batholith because it is intruded into marginal terranes that are not underlain by old continental crust. The region is also characterized by a widespread unconformity overlain by Eocene to Miocene volcanic and sedimentary strata, which are not present above the southern Sierra Nevada Batholith. Although Basin and Range normal faulting has disrupted the structural continuity of the northernmost Sierra Nevada Batholith, breaking it into tilted fault blocks, this has also exposed depth sections into the upper crust, supplying 3-D structural information on the upper crust of the batholith. Although the Sierra Nevada Batholith serves as a worldwide model for arc tectonics, the northern segment of this arc has received little study in the past. We address fundamental questions not thoroughly explored in the literature: Where exactly does the batholith go in the NW Basin and Range? How does the structure, age, petrology and geochemistry of the intrusions compare to those in the Sierra Nevada mountain range, and what can this tell us about the generation of arc batholiths? When did the batholithic roots of the magmatic arc become exposed at the surface, and where did the eroded material go?

What can this tell us about the paleogeography of the ancestral Sierra Nevada and its evolution to the modern range?

The basal Tertiary unconformity provides an important datum for reconstructing the pre-extensional geology of the area. Restored for Neogene extension, the geology v beneath this unconformity was compiled to construct an early Tertiary paleogeologic map. This map outlines the extent of granitic rocks of the Sierra Nevada Batholith and its country rocks across the NW Basin and Range. The paleogeologic map also illustrates the rock types eroded across the region prior to the development of the Tertiary unconformity and suggests deeper exhumation along the axis of the arc (compared to further east) in the late Cretaceous to early Tertiary.

To understand the structural and petrological architecture and history of this portion of the arc, eight 7.5´ quadrangles covering a distinct intrusive suite in the

Sahwave and Nightingale Ranges were mapped at 1:24,000 scale. Together, the

Sahwave and Nightingale ranges form a normal-fault-bound horst block that is warped into a gentle syncline. This is expressed in the dips of Miocene volcanic and sedimentary strata that unconformably overlie the Mesozoic batholith and its wallrocks. These basement rocks are composed of concentrically arranged granodiorite intrusions with an area of ~1000 km2, referred to as the Sahwave Intrusive Suite, intruded into older plutonic and metasedimentary rocks. The Sahwave Intrusive Suite youngs inward from equigranular hornblende biotite granodiorite to more felsic, K-feldspar megacrystic granodiorite. Zircon U-Pb dating by SHRIMP shows that the Sahwave Intrusive Suite spans ~ 93–88.5 Ma, and intrudes older plutonic rocks crystallized at 97–110 Ma.

Sample transects analyzed for modal mineralogy, trace and major element geochemistry show variation similar to that measured in coeval large zoned intrusions along the Sierra

Nevada crest, such as the Tuolumne Intrusive Suite, which represent the last and biggest pulse of magmatism in the Mesozoic arc. One substantial difference between the

Sahwave Intrusive Suite and coeval Sierran rocks are the more primitive Sr and Nd vi

87 86 isotopic ratios of the Sahwave Intrusive Suite ( Sr/ Sri ~ 0.7047 and εNd ~ –0.2). This is consistent with the inference that part of the arc in the central and southern Sierra Nevada straddles the edge of continental crust, whereas part of the arc in NW Nevada, which intrudes deep marine strata, is underlain by transitional or oceanic crust. The similarity of the Sahwave Intrusive Suite to its southern counterparts, despite this important difference, suggests arc flare-up events, as represented by the ca. 90 Ma intrusive suites along the Sierra Nevada crest, are controlled by a subcrustal process. Therefore, high magmatic flux might be caused by fluid release from the downgoing slab and/or an increase in subducted sediment rather than by backarc crustal thickening.

To constrain the detailed history of exhumation of the northern Sierra Nevada

Batholith, nine transects of samples were collected for thermochronologic study. Across the NW Basin and Range, normal faults tilt the batholithic rocks, exposing material that once lay up to 7 km depth beneath the early Tertiary unconformity. Apatite and zircon

U-Th/He, apatite fission track, biotite 40Ar/39Ar, and K-feldspar 40Ar/39Ar multi- diffusion-domain modeling were used together to understand the thermal history of the northern Sierra Nevada Batholith. Cooling and exhumation began promptly after intrusion in the Late Cretaceous, and continued at a roughly exponentially decreasing rate into the mid-Tertiary, consistent with previous work in the NW Basin and Range as well as the cooling history determined for the northern Sierra Nevada proper. The timing of exhumation is correlated with the local intrusive age of the batholith, and continues until later in the younger, eastern part of the arc. Correlations between magmatism and later exhumation suggest that residual arc heat acted to focus topographic relief and denudation. vii

ACKNOWLEDGEMENTS

Many people have supported me during the course of this research. Primary thanks go to my advisor, Elizabeth Miller: a constant source of no-nonsense geologic logic; excellent advice on which way to turn next; enthusiastic, foot-stamping motivation; and an appreciation for art and beauty in science. I thank Trevor Dumitru for training me in the intricacies of fission track dating and his finely honed mineral separation procedures. Joe Wooden has been an excellent mentor, both in the use and interpretation of ion microprobe data and also in introducing me to the visitors and new ideas always streaming through the SHRIMP lab. Gail Mahood has helped keep me honest on the workings of magmatic systems, and Norm Sleep has given me good advice on matters of heat flow. Marty Grove deserves credit for directing me in the use of 39Ar/40Ar dating and diffusion-domain modeling in K-feldspar, Jeremy Hourigan for help with U-Th/He dating at UC Santa Cruz, Bettina Wiegand for help measuring Sr and Nd isotopic data, and Bob Jones and Brad Ito for other lab help. I have also had fruitful collaborations with Sandra Wyld and Jim Wright, as well as excellent advice from Chris Henry and other folks at the Nevada Bureau of Mines and Geology. Special thanks go to Matt

Coble, true master of the argon line, and to Julie Fosdick for teaching me bomb digestion tactics for zircon. Thanks also to Joe Colgan for sharing data and samples from NW

Nevada, to Carrie Whitehill for getting me interested in the Sahwave and Nightingale

Ranges, and to the excellent companionship of many other grad students. Thanks to my field assistants Spencer Craven (2007), Laainam “Best” Chaipornkaew (2008), Silas

Stafford (2008), and Sarah Dasher (2009). Family and friends have provided unwavering support throughout this lengthy process. In particular I’d like to thank my wife, Sandy viii

Ottensmann, for everything from financial support to emotional support, and even domestic support when my working hours lengthened considerably.

This research was funded by two Stanford McGee grants, two Shell grants for travel, a GSA grad student research grant, and the U.S.G.S. National Cooperative

Geologic Mapping Program (EDMAP), as well as NSF Tectonics EAR-9417937, EAR-

9725371, EAR-0229854; Earthscope NSF 0346245; ACS-PRF 39063-AC8 and NSF

Tectonics 0809226. My tuition and stipend were supported by the Burt and Deedee

McMurtry Stanford Graduate Fellowship for three years, by the Carl H. Beal Memorial

Fund for one year, and by the Stanford University School of Earth Sciences for the remaining time.

This dissertation is dedicated to Brother Robert McDermott, S.J., (1916–2000), who gave me an early start on my career path as moderator of the Geology Paleontology

Society of Loyola High School. ix

TABLE OF CONTENTS

ABSTRACT ...... iv

ACKNOWLEDGEMENTS ...... vii

TABLE OF CONTENTS ...... ix

LIST OF TABLES ...... xii

LIST OF ILLUSTRATIONS ...... xiii

LIST OF PLATES ...... xiv

INTRODUCTION ...... 1

CHAPTER 1: EARLY TERTIARY PALEOGEOLOGIC MAP OF THE NORTHERN SIERRA NEVADA BATHOLITH AND THE NORTHWESTERN BASIN AND RANGE ...... 6

Abstract ...... 7 Introduction ...... 7 The basal Tertiary unconformity ...... 13 Paleogeologic map ...... 14 Discussion and conclusions ...... 15 Acknowledgements ...... 19 References ...... 19

CHAPTER 2: GEOLOGIC MAPPING OF THE SAHWAVE AND NIGHTINGALE RANGES, PERSHING COUNTY, NEVADA ...... 23

Abstract ...... 24 Introduction ...... 24 Geography of the study area ...... 26 Previous work ...... 29 Geologic setting ...... 30 Methods ...... 34 Lithologic units ...... 35 Triassic/Jurassic metasedimentary rocks ...... 36 x

Early intrusives ...... 37 Power Line Intrusive Complex ...... 37 Selenite Granodiorite ...... 38 Sahwave Intrusive Suite ...... 38 Granodiorite of Juniper Pass ...... 39 Granodiorite of Bob Spring ...... 40 Sahwave Granodiorite ...... 41 School Bus Granodiorite ...... 42 Minor intrusives ...... 42 Tertiary dikes ...... 44 Tertiary strata ...... 45 Quaternary units ...... 46 Structural development ...... 48 Jurassic to mid-Cretaceous deformation ...... 48 Emplacement of the Sahwave Intrusive Suite ...... 49 Mid-Miocene to present normal faulting ...... 52 Summary ...... 54 Acknowledgements ...... 55 References ...... 55

CHAPTER 3: THE SAHWAVE BATHOLITH, NW NEVADA: CRETACEOUS ARC FLARE-UP IN A BASINAL TERRANE ...... 62

Abstract ...... 63 Introduction ...... 64 Regional geologic setting ...... 67 Cretaceous plutonic rocks of northwest Nevada ...... 70 New mapping in the Sahwave and Nightingale Ranges ...... 74 Cenozoic strata ...... 75 Country rocks ...... 75 Early intrusive units ...... 80 The Sahwave Intrusive Suite ...... 81 xi

Minor intrusives ...... 84 Intrusive contacts ...... 85 Chronology of emplacement ...... 86 U-Pb SHRIMP methods ...... 87 U-Pb SHRIMP results ...... 87 Continuity of the Cretaceous Cordilleran batholith ...... 92 Mineralogy and geochemistry of the Sahwave Intrusive Suite ...... 94 Methods ...... 97 Mineralogy results ...... 100 Geochemistry results ...... 103 A Cathedral Range intrusive event outside the Sierra Nevada ...... 111 Emplacement ...... 115 Basinal setting and implications for arc flare-up ...... 118 Conclusions ...... 120 Acknowledgements ...... 121 References ...... 122 Appendix A: U-Pb SHRIMP analytical data ...... 132

CHAPTER 4: A MULTI-THERMOCHRONOMETER RECORD OF EROSIONAL EXHUMATION OF THE SIERRA NEVADA BATHOLITH IN THE NORTHWEST BASIN AND RANGE ...... 134

Abstract ...... 135 Introduction ...... 135 Geologic background ...... 143 Using thermochronology to constrain exhumation ...... 152 Sampling strategy ...... 156 Thermochronology results ...... 161 A. Pine Forest Range ...... 169 B. Santa Rosa Range ...... 170 C. Bloody Run Hills ...... 173 D. Seven Troughs Range ...... 174 E. Black Rock Range ...... 176 xii

F. Diamond Mountains ...... 177 G. Selenite Range ...... 179 H. Sahwave Range ...... 180 I. Wassuk Range ...... 184 Interpretation of the thermal histories ...... 186 Geothermal gradient ...... 186 Variations in exhumation among the transects ...... 190 Discussion ...... 193 Late Cretaceous to mid-Tertiary paleogeography ...... 193 Potential causes of erosional exhumation ...... 196 Summary ...... 200 Acknowledgements ...... 201 References ...... 202 Appendix B: Apatite fission track analytical procedures ...... 215 Appendix C: Apatite and zircon U-Th/He analytical procedures ...... 223 Appendix D: Biotite and K-feldspar 40Ar/39Ar analytical procedures ...... 230 Appendix E: U-Pb SHRIMP analytical procedures ...... 246

LIST OF TABLES Table 3.1. SHRIMP U-Pb geochronology sample data ...... 88 Table 3.2. Mineralogy and major element chemistry ...... 95 Table 3.3. Trace element chemistry ...... 98 Table 3.4. Sr and Nd isotopic data ...... 110 Table A1. U-Pb SHRIMP analytical data [Chapter 3] ...... 132 Table 4.1. Geological comparison of sample transects ...... 149 Table 4.2. Sample locality and age data ...... 159 Table B1. Apatite fission track analytical data ...... 218 Table C1. Zircon U-Th/He analytical data ...... 227 Table C2. Apatite U-Th/He analytical data ...... 228 Table D1. Biotite 40Ar/39Ar analytical data ...... 235 Table D2. K-feldspar 40Ar/39Ar analytical data ...... 236 xiii

Table E1. U-Pb SHRIMP analytical data [Chapter 4] ...... 247

LIST OF ILLUSTRATIONS Figure 1.1. Simplified geologic map of the NW Basin and Range ...... 8 Figure 1.2. Tertiary unconformity subcrop map of the NW Basin and Range ...... 10 Figure 1.3. Estimated extension map of the NW Basin and Range ...... 11 Figure 1.4. Restored paleogeologic map of the NW Basin and Range ...... 12 Figure 2.1. Simplified geologic map of the Sierra Nevada Batholith ...... 25 Figure 2.2. Shaded relief map of the Sahwave and Nightingale Ranges ...... 27 Figure 2.3. Photograph of the northern Sahwave Range ...... 28 Figure 2.4. Early Mesozoic terranes of the NW Basin and Range ...... 31 Figure 3.1. Shaded relief map of the western U.S. showing Mesozoic arc rocks ...... 65 Figure 3.2. Simplified geologic map of the northern Sierra Nevada Batholith ...... 66 Figure 3.3. Geologic map and cross section of the Sahwave Itrusive Suite ...... 72 Figure 3.4. Micrographs and field photographs of the rock units ...... 77 Figure 3.5. Zircon U-Pb SHRIMP results and concordia ...... 89 Figure 3.6. Cathodoluminescence images of representative zircons ...... 90 Figure 3.7. QAP diagram showing mineralogy of the Sahwave Intrusive Suite ...... 101 Figure 3.8. Radial variation in mineralogy ...... 102 Figure 3.9. Harker variation diagrams for selected major and trace elements ...... 104 Figure 3.10. Sahwave Intrusive Suite major element transect ...... 105 Figure 3.11. Rare earth element abundance ...... 107 Figure 3.12. Trace element abundance ...... 109 Figure 3.13. Sr and Nd isotopic variation ...... 112 Figure 3.14. Quartz-albite-orthoclase system: measured aplite compositions and minimum melt relations ...... 112 Figure 3.15. Simplified geologic map showing the Sierra Nevada Batholith and Cathedral Range–type intrusions ...... 113 Figure 3.16. Crustal-scale cartoon ...... 116 Figure 4.1. Map of the Cordilleran batholith in the western United States ...... 136 Figure 4.2. Map of early Tertiary unconformity and subcrop geology ...... 138 xiv

Figure 4.3. Map of NW Basin and Range showing thermochronology locations ...... 142 Figure 4.4. Geologic maps of the nine sampled ranges ...... 147 Figure 4.5. Geologic cross sections of the nine sampled ranges ...... 150 Figure 4.6. Zircon SHRIMP U-Pb results and concordia ...... 157 Figure 4.7. Thermochronology data from the nine transects ...... 162 Figure 4.8. Comparison of apatite fission track models from each transect ...... 167 Figure 4.9. Apatite fission track ages from the Sahwave Range ...... 183 Figure 4.10. Comparison of exhumation progress among the transects ...... 189 Figure 4.11. Geologic map of the study area showing the age of batholithic rocks and shallow fission track samples ...... 191 Figure 4.12. Schematic diagrams of exhumation across the arc ...... 194 Figure 4.13. Cartoons depicting possible exhumation scenarios ...... 198 Figure B1. Apatite fission track radial age plots ...... 219 Figure B2. Apatite fission track length distributions ...... 220 Figure B3. Apatite fission track modeling results ...... 221 Figure C1. Apatite (U-Th)/He ages and fitted activation energies ...... 229 Figure D1. Arrhenius plots showing K-feldspar argon diffusion results ...... 244

Figure D2. K-feldspar age and log(r/r0) spectra ...... 245

LIST OF PLATES Plate 1. Geologic Map of the Jayhawk Well Quadrangle ...... supplementary file Plate 2. Geologic Map of the Juniper Pass Quadrangle ...... supplementary file Plate 3. Geologic Map of the Tohakum Peak NE Quadrangle ...... supplementary file Plate 4. Geologic Map of the Tunnel Spring Quadrangle ...... supplementary file Plate 5. Geologic Map of the Bob Spring Quadrangle ...... supplementary file Plate 6. Geologic Map of the Tohakum Peak SE Quadrangle ...... supplementary file Plate 7. Geologic Map of the Sage Hen Spring Quadrangle ...... supplementary file Plate 8. Geologic Map of the Bob Spring Quadrangle ...... supplementary file

INTRODUCTION

The Sierra Nevada Batholith is among the most thoroughly studied of all magmatic arcs. Formed during the Mesozoic by eastward subduction of oceanic crust beneath North American crust and previously accreted terranes, the Sierra Nevada

Batholith has guided decades of thought about subduction tectonics and plutonism in general. Most studies have focused on the beautiful, glaciated outcrops of the high Sierra

Nevada. However, plutonic rocks of similar character and age, though buried and disrupted by Tertiary volcanism and extension, continue well north of the oft-studied outcrops, past the limits of the Sierra Nevada mountains, past the edge of continental crust, and past the limits of many published figures, into the desolate, sage-covered wilds of northwest Nevada. This northernmost part of the Sierra Nevada Batholith forms the subject of my dissertation.

Some recent studies have begun to challenge long-held notions about the Sierra

Nevada Batholith, even calling into question whether the batholith is a good archetype for arc magmatism. For example, some posit that the batholith’s Late Cretaceous flare-up is related to the underthrusting of crustal material from a backarc orogenic plateau, the

“Nevadaplano” (Ducea, 2001; DeCelles and Coogan, 2006; Decelles et al., 2009). Others have interrogated the history of topographic development of the Sierra Nevada with thermochronometry and stable isotope data, suggesting the high elevations of the range are not recently generated, as previously thought (e.g., Whitney, 1880; Lindgren, 1911;

Wakabayashi and Sawyer, 2001), but are a mere relict of the Cretaceous ancestral Sierra

Nevada or the rim of the Nevadaplano (e.g., House et al., 1998, 2001; Mulch et al., 2006;

Cassel et al., 2009). More controversial hypotheses ask whether differentiation of the 2 batholith might have been the result of zone refining in a thermal gradient rather than crystal-liquid fractionation (Lundstrom, 2009), whether Sierran plutons were the roots of a substantial volcanic arc (e.g., Glazner et al., 2008), or whether the arc was created by east-dipping subduction at all (Hildebrand, 2009).

The northernmost Sierra Nevada Batholith is a bit different from the rest and is therefore a good place to address some of these controversies from a new point of view.

The most notable differences are (1) intrusion into a basinal terrane rather than along the edge of true continental crust, (2) significant disruption and 3-D exposure of the batholith by Neogene extensional faulting, and (3) the incision of the batholith by a widespread early Tertiary (pre-extensional) unconformity. The first difference allows us to test the influence of crustal composition on arc magmatism. Additionally, the basinal Triassic terrane of NW Nevada does not appear to have ever been underlain by overthickened crust of the sort expected beneath a major orogenic plateau (Colgan et al., 2006; Lerch et al., 2007). This makes it an ideal testing ground for models that link parts of the intrusion and uplift history of the Sierra Nevada Batholith to the tectonics of the Nevadaplano.

The effect of Basin and Range extension is more of a mixed blessing. On the one hand, normal faulting resulted in disruption of basement structures, extensive cover by

Cenozoic basin-fill deposits, and a dry climate without the history of glacial erosion that makes such spectacular granitic outcrops in the Sierra Nevada mountains. On the other hand, normal faulting dissects the batholith and exposes depth sections up to ~ 7 km below the level of the basal Tertiary unconformity (e.g., Stockli et al., 2002), giving a unique 3-D view into upper arc crust. The position of the Mesozoic arc in the Basin and

Range is also important for evaluating the relationship between the modern Sierra 3

Nevada mountains and their namesake batholith, in light of hypotheses that link the two closely (e.g., Jones et al., 2004). The presence of a regional basal Tertiary unconformity provides a datum useful for understanding the distortions caused by Neogene extensional structures and for restoring the basement geology to its pre-extensional configuration.

Because so little previous data exist for the northern end of the Sierra Nevada

Batholith, much of this dissertation is necessarily aimed at laying the groundwork for future study—complete resolution of the Sierra Nevada Batholith’s tectonic history is beyond the scope of this work. Nevertheless, I address some basic questions about the paleogeology and paleogeography of the region; I use detailed mapping, petrological, and geochemical data from a specific intrusive suite in NW Nevada to understand how arc magmatism developed across the basinal terrane and how that compares with batholith development along the edge of continental crust; and I use a variety of thermochronological data from normal-fault-exposed depth transects into the batholith to explore its Late Cretaceous to mid-Tertiary erosional exhumation. In short, I focus on the tectonic evolution of the northern Sierra Nevada Batholith from the time of its Cretaceous emplacement until its erosional history, until net deposition of strata above the regional unconformity began in the Eocene to Miocene.

Chapter 1 presents a compilation of subcrop data from the Tertiary unconformity, which is restored into an early Tertiary paleogeologic map of the northern Sierra Nevada

Batholith and adjacent parts of the Basin and Range. This map is used to clarify the extent of the Sierra Nevada batholith in the Basin and Range, to estimate the magnitude of Late Cretaceous/early Tertiary erosional exhumation of the arc (which can be used as a proxy to its uplift), and to constrain what basement rock types and ages could have 4 contributed to early Tertiary detritus carried out of the region by rivers and deposited elsewhere. This work is published in Geology (Van Buer et al., 2009), with Elizabeth

Miller and Trevor Dumitru as coauthors.

Chapter 2 presents field data, 1:24,000 scale geologic mapping, and cross sections of eight 7.5´ quadrangles in the Sahwave and Nightingale Ranges of northwest Nevada, encompassing a previously unknown, large (~1000 km2), concentrically zoned intrusion, referred to as the Sahwave Intrusive Suite. Lithologic descriptions and a summary of geologic events represented in the maps and cross sections (Plates 1–8) elucidate the context and detailed structural and petrologic architecture of one of the best-exposed intrusions of the batholith in NW Nevada. A version of these maps, reduced to 1:62,500 scale, has been submitted for publication as a Nevada Bureau of Mines and Geology

Open-File Report.

Chapter 3 presents geochronology, modal mineralogy, and geochemistry for the

Sahwave Intrusive Suite, in the context of relevant field data distilled from Chapter 2.

Zircon U-Pb SHRIMP ages reveal a history of magmatism beginning at ca. 110 Ma and culminating with intrusion of the volumetrically dominant ~ 93–88.5 Ma Sahwave

Intrusive Suite, similar to the Late Cretaceous, eastern part of the arc in the Sierra

Nevada. The petrology and geochemistry of the Sahwave Intrusive Suite are compared to the Tuolumne Intrusive Suite and contemporaneous high-flux intrusions along the Sierran crest. Strong similarities, despite a very different crustal context, suggest a subcrustal mechanism controls arc flare-up. A version of this chapter was published in Lithosphere

(Van Buer and Miller, 2010) with Elizabeth Miller as coauthor. 5

Chapter 4 addresses the erosional exhumation of the northern Sierra Nevada batholith up to the time of deposition above the early Tertiary unconformity using data collected with multiple thermochronometers from nine fault-exposed transects in the NW

Basin and Range Province. Temperature-time information from apatite and zircon U-

Th/He, apatite fission track, and 40Ar/39Ar K-feldspar multi-diffusion-domain modeling are interpreted in the context of their thermal evolution as a function of depth beneath the

Tertiary unconformity to create a detailed cooling history for this part of the northern

Sierra Nevada Batholith. Erosional exhumation began promptly after intrusion of the batholith ceased in the Late Cretaceous but continued, at a decreasing rate, into the mid-

Tertiary, with substantial rock uplift continuing until a later time in the east than in the west. A version of this chapter is in preparation for submission under the authorship of

Nicholas Van Buer, Trevor Dumitru, Marty Grove, and Jeremy Hourigan. 6

CHAPTER 1: EARLY TERTIARY PALEOGEOLOGIC MAP OF THE NORTHERN SIERRA NEVADA BATHOLITH AND THE NORTHWESTERN BASIN AND RANGE

Nicholas J. Van Buer, Elizabeth L. Miller and Trevor A. Dumitru Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305

A version of this chapter has been published in Geology and copyright has been assigned to the Geological Society of America. GSA’s copyright policy explicitly permits reproduction in an author’s dissertation.

2009 Geological Society of America Van Buer, N.J., Miller, E.L., Dumitru, T.A., 2009, Early Tertiary paleogeologic map of the northern Sierra Nevada batholith and the northwestern Basin and Range: Geology, v. 37, p. 371-374, doi: 10.1130/G25448A.1. 7

ABSTRACT

The northern Sierra Nevada and adjacent Basin and Range are marked by a widespread regional unconformity at the base of the Eocene to Miocene volcanic and sedimentary section that overlies the Mesozoic batholith and its wallrocks. To help address controversial questions about the origin, uplift, and erosion of the batholith, we compiled a subcrop geologic map of the unconformity prior to Tertiary extension. This simple but underutilized technique reveals the distribution of rock units that could have contributed detritus to the Eocene-Miocene river channels crossing the Sierra Nevada and demonstrates that the Mesozoic Sierra Nevada batholith was continuous to the northeast across the NW Basin and Range. More speculatively, the subcrop map implies that Late

Cretaceous to Eocene erosional stripping may have been greatest above the axis of the batholith and decreased to the east; thus, drainage in this area may have been eastward and then switched westward in Eocene to Miocene times.

INTRODUCTION

An extensive regional unconformity, termed the "basal Tertiary unconformity", is currently exposed at higher elevations in the northern part of the Sierra Nevada and in numerous mountain ranges across much of the northern Basin and Range province. This unconformity separates underlying late Mesozoic plutonic and country rocks of the Sierra

Nevada Batholith from immediately overlying Eocene to Miocene stratified volcanic and sedimentary rocks (Fig. 1.1). As such, the unconformity represents the Earth's surface in this region at the time of the fundamental tectonic transition from uplift and erosion of the extinct Sierra Nevada magmatic arc to deposition and preservation of the overlying 8

42° Yreka 1.1 N M O D O C N P L A T E A U KLAMATH RANGE 100 km

NEVADA

CALIFORNIA

S I E R R A N E V A D A B A S I N A N D R A N G E

38° 121° W 116° W N

Legend

Cenozoic cover

Granitic rocks of the Sierra Nevada Batholith

Other Mesozoic and Paleozoic rocks

Figure 1.1. Plutonic rocks of the northern Sierra Nevada batholith (red) and their country rocks (green) extend across the northern Basin and Range but are dismembered by normal faulting and partly covered by Tertiary strata. 9 strata. In this paper, we compile a regional map of the rock units present immediately beneath the unconformity (Fig. 1.2) and then remove post-unconformity Basin and Range extension (Fig. 1.3). This yields a paleogeologic map of the region's land surface in early to middle Tertiary time (Fig. 1.4).

Our reconstruction of the pre-extensional surface geology across a broad region provides useful and incontrovertible data bearing on several debated topics regarding the tectonic and topographic evolution of this region. It clarifies the pre-Tertiary extent and geology of the northern part of the Sierra Nevada Batholith prior to Basin and Range faulting (Fig. 1.1). It also helps address the pre-extensional topography of the region by illustrating variations in depth of erosion beneath the unconformity. Estimates of batholith and back-arc crustal thickness based on restoration of regional cross sections

(e.g., Chase et al., 1998; DeCelles et al., 1995), mineral thermobarometry (e.g., Camilleri et al., 1997; Lewis et al., 1999), and paleotopography using oxygen isotopes (e.g., Mulch et al., 2006) all suggest that thick crust (50-65 km) and high elevations (2-4 km) persisted to the beginning of Tertiary extension. Conversely, restoration of known amounts of

Tertiary extension together with geophysical data on Moho depths in NW Nevada suggest “normal” (~38 km) thickness crust prior to extension (e.g., Colgan et al., 2006;

Lerch et al., 2007). Analysis of our early Tertiary paleogeologic map represents an independent approach to addressing the controversial paleogeology and paleotopography across the region. Although the construction of such maps has proven extremely useful for understanding pre-Basin and Range structure in the Sevier thrust belt (Armstrong,

1968) and in east-central Nevada (Gans and Miller, 1983), until now it has not been 10

42° N 1.2 2004 Stanford seismic line (Lerch et al., 2007) Mzg Owyhee Plateau colored by inferred basement (no data)

Modoc Plateau (no data) lPz Elko Mzm Mzm PMs H uPz uPz uPz P Mzg

No data

lPz T Mzm No data lPz N

uPz uPz Mzg Mzm 100 km 38° N 121° W 116° W Legend Unconformity (color coded by basement)

Mzg Granitic rocks of the Sierra Nevada Batholith

Mzm Mesozoic wall rocks of the batholith

uPz Upper Paleozoic wall rocks and Golconda allochthon

lPz Lower Paleozoic wall rocks and Roberts Mts. allochthon

PMs Paleozoic to Mesozoic miogeoclinal sequence (anticlines expose rocks as old as Ordovician; synclines as young as Jurassic) Figure 1.2. Data base of the extent and nature of the basal Tertiary unconformity and underlying rocks across northern Nevada and the northern Sierra. Line squiggles represent exposed traces of the unconformity today; color-coding represents the main rock types immediately beneath the unconformity. Colored fields represent a post-extensional view of the pre-unconformity paleogeologic map. H = Honey Lake, P = Pyramid Lake; T = Lake Tahoe. 11

121° W 116° W 1.3 Owyhee Plateau Modoc Plateau 42° N <5% since 5-10%, @ 0-14 Ma NW Nevada 16 Ma 15-25% (0%) (7.5%) @ 0-12 Ma NE Nevada (20%) 15-20% @ 0-12 Ma older Eocene? (20%) N

100 km Central Nevada 15-100% @ 16 Ma, older Oligocene (50%) WLZ Sierra Front 5-10% @0-10? Ma (7.5%) 38° N Yerington >>100% @14 Ma, 5-10% @0-5 Ma (150%)

Figure 1.3. Generalized estimates of percent extension and its age east of the Sierra Nevada, based on fault attitudes, tilt of the Tertiary unconformity, and palinspastic restoration of cross sections. NW Nevada – Colgan et al. (2006); Singatse-Wassuk Range – Surpless et al. (2002) and Proffett and Dilles (1984); central Nevada –Smith et al. (1991); Modoc-Owyhee Plateau – Wells and Heller (1988). Actual values used for restoration shown in parentheses.

Figure 1.4 (next page): Restored early Tertiary paleogeologic map with extension removed. LFTB = Luning-Fencemaker Thrust Belt, GT = Golconda Thrust, RMT = Roberts Mountains Thrust. The graph at bottom shows the relative amounts of erosion at the batholith and back-arc region inferred along a schematic west to east cross section along line A-A’. 12

1.4

▼

▲

▼

A’

▲

▼

▲

▼

▲

▼

▲ ▼

▼

▲

▼

▲

▼

▲

▼

▲ A ▼

▼

▲

▼

RMT

▼

▲

LFTB▼

▼

▲

▼

▼ ▲

▲

▼

▲

▼

▼ ▲

▼

▲ ▼

▲ 100 km

5 km Approximate depth of erosion 0 A A’

Same sacle as above map; no vertical exaggeration

Cenozoic cover (no data)

▼

Granitic rocks of the Sierra Nevada Batholith

▼ Paleozoic folds and thrusts

▼

Mesozoic wall rocks of the batholith

▼ Mesozoic folds and thrusts Upper Paleozoic and Golconda allochthon

Lower Paleozoic and Roberts Mts. allochthon Paleo-river canyons

Paleozoic to Mesozoic miogeoclinal sequence Paleo-continental divide (anticlines expose rocks as old as Ordovician; (both from Henry, 2008) synclines as young as Jurassic) 13 comprehensively applied across the northern Sierra Nevada and adjacent Basin and

Range.

THE BASAL TERTIARY UNCONFORMITY

The striking unconformity at the base of the predominately volcanic Tertiary section in the northern Sierra Nevada forms a mostly unfaulted, continuous surface developed across granitoids and associated metamorphosed wallrocks (e. g., Lindgren,

1911) (Fig. 1.2). In the Basin and Range, younger extensional faulting offsets and tilts this same unconformity, exposing it in many individual fault-bound ranges (Fig. 1.2). In detail, the unconformity is somewhat diachronous, with basal overlying rocks varying in age from Lower Eocene (gold-bearing gravels in the Sierra paleochannels) to Miocene volcanic rocks in both the Sierra Nevada and Basin and Range. In the Basin and Range, southward younging (Eocene to Miocene) of the oldest volcanic rocks above the unconformity has been well documented (Stewart, 1998). The relief of the unconformity is generally quite moderate, but is as much as 0.5-1.6 km in deep paleochannels that preserve Eocene to Miocene fills in many places throughout our map area (Fig. 1.4)

(Henry, 2008). The limited stratigraphic thicknesses of Eocene to Miocene rocks suggests that deposition was minor during this interval, whereas the widespread preservation of these same rocks suggests that erosion was also minor during this time span compared to the deep erosion of the batholith between the end of magmatism at about 85 Ma and the early development of the unconformity at about 40 Ma.

PALEOGEOLOGIC MAP

The greater geologic complexity of the Basin and Range province relative to the

Sierra Nevada (Fig. 1.1) is primarily a function of Basin and Range normal faulting, which has exposed a wide range of crustal levels at the surface today. Prior to Basin and

Range extension, the map distribution and continuity of units would have been far simpler. To illustrate this, we first constructed the map in Figure 1.2 showing the modern surface traces of the basal Tertiary unconformity as wormy lines color-keyed to the subjacent rock units. We compiled this unconformity data at 1:500,000 scale from numerous local and county geologic maps. In areas with no unconformity exposure, but shallow pre-extensional paleodepths, we inferred the subjacent unit from modern map patterns. Only plutons > 10 km in diameter are shown.

To construct a paleogeologic map from Figure 1.2, the effects of Cenozoic normal faulting were removed, at least in an approximate way. Extension in the northern Basin and Range over the last 50 million years has been episodic and heterogeneous, with different regions extending at different times (and/or multiple times) and by different amounts (e.g., Miller et al., 1999; Stockli, 2000; Surpless et al., 2002; McQuarrie and

Wernicke, 2005; Colgan et al., 2006). Inception of major Basin and Range faulting began at 17 Ma in central Nevada, 12 Ma in northern Nevada and is believed to be younger than

8 Ma along the eastern Sierra Nevada (Fig. 1.3). A number of separate extensional domains have been defined, where percent extension has been estimated through the construction of balanced cross sections or consideration of Tertiary unconformity data

(Fig. 1.3). The strain domain map is highly generalized, but at the scale of Fig. 1.2 can be used to qualitatively restore the sub-unconformity units into their approximate position 15 prior to Cenozoic faulting (Fig. 1.4). The choice of percent extension was based on conservative estimates, and there are parts of the map where data are unavailable (e.g. the

Carson Sink). Because the total magnitude of extension is not large (0-50%, except for local regions, Fig. 1.3), minor changes to our assumed extensional domains have little effect on the resulting map. Extension was restored in the direction of fault slip, based on geologic mapping and/or perpendicular to the modern range-bounding normal faults for each domain (parallel to the arrows shown in Fig 1.3). Restoration also accounted for ~

25-60 km of right-lateral movement (decreasing to the north) across the Walker Lane

(Wesnousky, 2005; Faulds et al., 2005).

DISCUSSION AND CONCLUSIONS

Several fundamental relations are apparent from Figure 1.4. First, this map provides the framework necessary for understanding detrital signatures preserved in the

Eocene to Miocene channels of the Sierra Nevada. According to Figure 1.4, these channels should primarily contain detritus derived from the batholith and its overlying

Tertiary volcanic units, upper Paleozoic to lower Mesozoic strata of the miogeocline,

Paleozoic eugeoclinal rocks of the Roberts Mountains and Golconda allochthons, and weakly metamorphosed Mesozoic basinal rocks from western Nevada. If the paleodrainages were at some time different than depicted in Fig. 1.4, this should also be reflected in the detrital signatures. Regardless, channel deposits should not contain material derived from stratigraphically deeper rocks of the miogeocline or from the plutons that intrude these depths, and preliminary detrital zircon studies of Eocene channel deposits in the Sierra suggest this is the case (Cecil and Ducea, 2007). 16

Second, the continuity of the Sierra Nevada Batholith across NW Nevada becomes apparent after Cenozoic cover and extension are removed. Unlike the southern

Sierra Nevada Batholith, which trends NNW along the truncated edge of the North

American craton, parallel to the modern continental margin, the northern part of the batholith trends NNE and intrudes Mesozoic back-arc basin deposits, bending inland from the continental margin. The buried NW edge of the batholith (Fig. 1.4.) is imaged geophysically as a boundary between lower (to the east) and higher (to the west) seismic velocities in the upper 15 km of the crust (Fig. 1.2; Lerch et al., 2007). Petrologic and geochronologic data indicate that the granitic rocks crossing the NW Basin and Range correlate with the mid- to late-Cretaceous plutonic belts of the Sierra Nevada Batholith

(Smith et al., 1971; Barton et al., 1988; Lerch et al., 2007; Van Buer, 2008). The magnitude of Tertiary extension in the NW Basin and Range is well constrained (Colgan et al., 2006) and insufficient to account for the eastward bend in the Sierra Nevada

Batholith (Fig. 1.3, 1.4). This suggests that the bend is a primary feature of the

Cretaceous arc. Northward shallowing of the subduction angle and/or batholith-age or younger extension beneath the Modoc plateau (Nilsen, 1993) could explain the bend.

Third, it appears that more material was eroded from the axis of the batholith than from the back-arc region to the east during late Cretaceous to Eocene time (Fig. 1.4).

Nearly all traces of coeval volcanic cover were eroded from the northern Sierra Nevada

Batholith. Granitoid emplacement depths at the unconformity are 3-8 km (e.g., Ague,

1997). Narrow, andalusite-bearing contact aureoles associated with satellite plutons indicate < 7-11 km depths for the batholith in NW Nevada (Compton, 1960; Colgan

2006). Many Mesozoic plutons are currently exposed to the east of those shown in 17

Figure 1.2, but they intrude crustal levels that remained at depth during the development of the Tertiary unconformity and were only exposed by Basin and Range faulting and tilting (e.g. Smith et al., 1991; Gans and Miller, 1983). From the east side of the batholith to the Luning-Fencemaker thrust system (Figure 1.4), the metamorphic grade of deformed Mesozoic basinal deposits varies from low to medium greenschist facies (Wyld et al., 2003). East of the Luning-Fencemaker belt, the Paleozoic rocks of the Golconda and Roberts Mountains allochthons are mostly unmetamorphosed although locally overprinted by Mesozoic deformation and metamorphism (Miller et al., 1992). Mesozoic deformation of these older tectonic units is sufficiently limited at the regional scale that the Paleozoic thrust architecture appears fundamentally unmodified at the scale of our paleogeologic map (Figure 1.4). East of the Roberts Mountains thrust (Fig. 1.4), upper

Paleozoic and lower Mesozoic miogeoclinal strata beneath the Tertiary unconformity have experienced little or no burial-induced metamorphism, as determined by their conodont alteration indices (e.g., Gans et al., 1990). Sevier-age folds in the miogeocline east of the Roberts Mountains thrust expose rocks as deep as Ordovician in anticlines, but essentially unburied Jurassic strata are preserved in synclines, consistent with erosion varying from 0-4 km. Overall, these observations indicate that depths of post-batholith erosion beneath the Tertiary unconformity decrease eastward from the axis of the batholith.

Greater erosion along the batholith axis implies greater net uplift and, presumably, higher topography in this region generated during its emplacement history.

This relationship supports the idea that the ancestral Sierra Nevada formed a paleo-divide

(e.g. Wakabayashi and Sawyer, 2001; DeGraaff-Surpless et al., 2002), with rivers 18 draining both westwards, into the Great Valley, and also eastwards or northwards (Fig.

1.4). Modest pre-extensional crustal thicknesses (~38 km) in NW Nevada (e.g. Colgan et al., 2006) as compared to the southern Sierra Nevada (~45-65 km, e.g., Wernicke et al.,

1996), however, suggest that this northern topographic divide may not have been as high as the contemporaneous southern Sierra Nevada. By Eocene time, however, paleo-rivers draining westward from a divide in central/eastern Nevada had already been incised and were beginning to accumulate gravels and volcanogenic strata, as documented in detail by Henry (2008) (Fig. 1.4). The divide or crest of the western U.S. at this time, as shown by Henry (2008), is depicted as a dashed line in Fig. 1.3, and lies considerably to the east.

Where studied by Henry (2008), Eocene to Miocene river systems persisted until disruption by Basin and Range faulting and carried debris from as far east as central

Nevada across the Sierra (Figure 1.4). These relations suggest a major shift in paleogeography within the time interval bracketed by the youngest plutons of the batholith (~85 Ma) and the oldest fluvial deposits (~45 Ma) of the Eocene to Miocene river systems of the Sierra (Fig. 1.4) (e.g. Lindgren, 1911; Henry, 2008). This shift was followed by the major switch from erosion to deposition as marked by the unconformity.

Future work that better defines the timing of this regional paleotopographic change will be very useful for constraining geodynamic models of the change from batholith emplacement to shallow-slab subduction to renewed magmatism and the onset of Basin and Range extension.

ACKNOWLEDGMENTS

This paper represents part of an ongoing effort by the structural geology and tectonics research group at Stanford aimed at compiling and interpreting data on the

Tertiary unconformity across the northern Great Basin and Sierra Nevada. During this time-span, our research was partially funded by NSF Tectonics EAR-9417937, EAR-

9725371, EAR-0229854; Earthscope NSF 0346245; ACS-PRF 39063-AC8 and NSF

Tectonics 0809226.

REFERENCES

Ague, J.J., 1997, Thermodynamic calculation of emplacement pressures for batholithic rocks, California: Implications for the aluminum-in-hornblende barometer: Geology, v. 25, p. 563–566, doi: 10.1130/0091- 7613(1997)025<0563:TCOEPF>2.3.CO;2.

Armstrong, R.L., 1968, Sevier orogenic belt in Nevada and Utah: Geological Society of America Bulletin, v. 79, p. 429–458, doi: 10.1130/0016- 7606(1968)79[429:SOBINA]2.0.CO;2.

Barton, M.D., Battles, D.A., Debout, C.E., Capo, R.C., Christensen, J.N., Davis, S.R., Hanson, R.B., Michelson, C.J., and Trim, H.G., 1988, Mesozoic contact metamorphism in the western United States, in Ernst, W.G., editor, Metamorphism and crustal evolution of the western United States, Rubey Volume 7: Englewood Cliffs, New Jersey, Prentice Hall, p. 110–178.

Camilleri, P., Yonkee, A., Coogan, J., DeCelles, P., McGrew, A., and Wells, M., 1997, Hinterland to foreland transect through the Sevier Orogen, northeast Nevada to north central Utah; structural style, metamorphism, and kinematic history of a large contractional orogenic wedge: Geology Studies, v. 42, p. 297–309.

Cecil, M.R., and Ducea, M.N., 2007, Constraining landscape evolution in the Sierra Nevada through detrital zircon provenance analysis of Cenozoic fluvial sediments: Geological Society of America Abstracts with Programs, v. 39, no. 6, p. 436.

Chase, C.G., Gregory, K.M., Parrish, J.T., and DeCelles, P.G., 1998, Topographic history of the western Cordillera of North America and the etiology of climate, in Crowley, T.J., Burke, K., editors, Tectonic boundary conditions for climate 20

reconstructions: Oxford Monographs on Geology and Geophysics, v. 39, p. 73– 99.

Colgan, J.P., Dumitru, T.A., Reiners, P.W., Wooden, J.L., and Miller, E.L., 2006b, Cenozoic tectonic evolution of the Basin and Range Province in northwestern Nevada: American Journal of Science, v. 306, p. 616–654, doi: 10.2475/08.2006.02.

Compton, R.R., 1960, Contact metamorphism in the Santa Rosa Range, Nevada: Geological Society of America Bulletin, v. 71, p. 1383–1416, doi: 10.1130/0016- 7606(1960)71[1383:CMISRR]2.0.CO;2.

DeCelles, P.G., Lawton, T.F., and Mitra, G., 1995, Thrust timing, growth of structural culminations, and synorogenic sedimentation in the type Sevier orogenic belt, Western United States: Geology, v. 23, p. 699–702, doi: 10.1130/0091- 7613(1995)023<0699:TTGOSC>2.3.CO;2.

DeGraaff-Surpless, K., Graham, S.A., Wooden, J.L., and McWilliams, M.O., 2002, Detrital zircon provenance analysis of the Great Valley Group, California: Evolution on an arc-forearc system: Geological Society of America Bulletin, v. 114, p. 1564–1580, doi: 10.1130/0016- 7606(2002)114<1564:DZPAOT>2.0.CO;2.

Faulds, J.E., Henry, C.D., and Hinz, N.H., 2005, Kinematics of the northern Walker Lane: An incipient transform fault along the Pacific-North American plate boundary: Geology, v. 33, p. 505–508, doi: 10.1130/G21274.1.

Gans, P.B., Repetski, J.E., Harris, A.G., and Clark, D.H., 1990, Conodont geothermometry of Paleozoic supracrustal rocks in the eastern Great Basin, in Geology and Ore Deposits of the Great Basin; Programs with Abstracts, Cuffney, B., editor, Geological Society of Nevada, p.103.

Gans, P.B., and Miller, E.L., 1983, Field Trip 6; Style of mid-Tertiary extension in east- central Nevada: Special Studies – Utah Geological and Mineral Survey, v. 59, p. 107–160.

Glazner, A.F., Coleman, D.S., and Bartley, J.M., 2008, The tenuous connection between high-silica rhyolites and granodiorite plutons: Geology, v. 36, p. 183-186.

Henry, C.D., 2008, Ash-flow tuffs and paleovalleys in northeastern Nevada: Implications for Eocene paleogeography and extension in the Sevier hinterland, northern Great Basin: Geosphere, v. 4, p. 1–35, doi: 10.1130/GES00122.1.

Hildebrand, R.S., 2009, Did westward subduction cause Cretaceous-Tertiary orogeny in the North American Cordillera? Special Paper – Geological Society of America, v. 457, 71 pp. 21

Lerch, D.W., Klemperer, S.L., Glen, J.M.G., Ponce, D.A., and Miller, E.L., 2007, Crustal structure of the northwestern Basin and Range province and its transition to unextended volcanic plateaus. Geochemistry, Geophysics, Geosystems, v. 8 (1) doi:10.1029/2006GC001429, 21 p.

Lewis, C.J., Wernicke, B.R., Selverstone, J., and Bartley, J.M., 1999, Deep burial of the footwall of the northern Snake Range decollment, Nevada: Geological Society of America Bulletin, v. 111, p. 39–51, doi: 10.1130/0016- 7606(1999)111<0039:DBOTFO>2.3.CO;2.

Lindgren, W., 1911, The Tertiary gravels of the Sierra Nevada of California, U. S. Geological Survey Professional Paper 73, 226 p.

Lundstrom, C.C., 2009, Hypothesis for the origin of convergent margin granitoids and Earth’s continental crust by thermal migration zone refining: Geochimica et Cosmochimica Acta, v. 73, p. 5709-5729.

McQuarrie, N., and Wernicke, B.P., 2005, An animated tectonic reconstruction of southwestern North America since 36 Ma: Geosphere, v. 1, p. 147-172.

Miller, E.L., Miller, M.M., Wright, J.E., and Madrid, R., 1992, Late Paleozoic paleogeographic and tectonic evolution of the Western U.S. Cordillera, in The Cordilleran Orogen; Conterminous U.S., Burchfiel, B.C., Lipman, P.W., and Zoback, M.L., editors, DNAG, p. 57–106.

Miller, E.L., Dumitru, T.A., Brown, R.W., and Gans, P.B., 1999, Rapid Miocene slip on the Snake Range-Deep Creek Range fault system, east-central Nevada: Geological Society of America Bulletin, v. 111, p. 886–905, doi: 10.1130/0016- 7606(1999)111<0886:RMSOTS>2.3.CO;2.

Mulch, A., Graham, S.A., and Chamberlain, C.P., 2006, Hydrogen isotopes in Eocene river gravels and paleoelevation of the Sierra Nevada: Science, v. 313, p. 87–90, doi: 10.1126/science.1125986.

Nilsen, T., 1993, Stratigraphy of the Cretaceous Hornbrook Formation, southern Oregon and northern California, U. S. Geol. Surv. Prof. Pap., P1521.

Proffett, J.M., Jr., and Dilles, J.H., 1984, Geologic map of the Yerington District, Nevada: Nevada Bureau of Mines and Geology Map 77, scale 1:24 000, 1 sheet.

Smith, D.L., Gans, P.B., and Miller, E.L., 1991, Palinspastic restoration of Cenozoic extension in the central and eastern Basin and Range Province at latitude 39–40 degrees N, in Geology and ore deposits of the Great Basin; symposium proceedings, Raines, G.L., Lisle, R.E., Schafer, R.W., and Wilkinson, W.H., editors, Geological Society of Nevada, Reno, p. 75–86. 22

Smith, J.G., McKee, E.H., Tatlock, D.B., and Marvin, R.F., 1971, Mesozoic granitic rocks in northwestern Nevada: A link between the Sierra Nevada and Idaho Batholiths: Geological Society of America Bulletin, v. 82, p. 2933–2944, doi: 10.1130/0016-7606(1971)82[2933:MGRINN]2.0.CO;2.

Stewart, J.H., 1998, Regional characteristics, tit domains, and extensional history of the later Cenozoic Basin and Range Province, western North America: Geological Society of America Special Paper, v. 323, p 47–74.

Stockli, D.F., 2000, Regional timing and spatial distribution of Miocene extension in the northern Basin and Range [Ph.D. Thesis]: Stanford University, 239 p.

Surpless, B.E., Stockli, D.F., Dumitru, T.A., and Miller, E.L., 2002, Two-phase westward encroachment of Basin and Range extension into the northern Sierra Nevada: Tectonics, v. 21, doi: 10.1029/2000TC001257.

Van Buer, N.J., 2008, The Cathedral Range intrusive event in western Nevada: new constraints: Geological Society of America Abstracts with Programs, v. 40, no. 1, p. 92.

Wakabayashi, J., and Sawyer, T.L., 2001, Stream incision, tectonics, uplift, and evolution of topography of the Sierra Nevada, California, Journal of Geology, v. 109 (5), p. 539-562.

Wells, R.E., and Heller, P.L., 1988, The relative contribution of accretion, shear, and extension to Cenozoic tectonic rotation in the Pacific Northwest: Geological Society of America Bulletin, v. 100, p. 325–338, doi: 10.1130/0016- 7606(1988)100<0325:TRCOAS>2.3.CO;2.

Wernicke, B.P. and eighteen others, 1996, Origin of high mountains in the continents; the southern Sierra Nevada: Science, v. 271, p. 190-193.

Wesnousky, S.G., 2005, The San Andreas and Walker Lane fault systems, western North America: transpression, transtension, cumulative slip and the structural evolution of a major transform plate boundary: Journal of Structural Geology, v. 27, p. 1505–1512, doi: 10.1016/j.jsg.2005.01.015.

Whitney, J.D., 1880, The auriferous gravels of the Sierra Nevada of California: Harvard University, Museum of Comparative Zoology, Memoirs, v. 1, 659 pp.

Wyld, S.J., Rogers, J.W., and Copeland, P., 2003, Metamorphic evolution of the Luning- Fencemaker fold-thrust belt, Nevada: Illite crystallinity, metamorphic petrology, and 40Ar/39Ar geochronology: The Journal of Geology, v. 111, p. 17–38, doi: 10.1086/344663.

CHAPTER 2: GEOLOGIC MAPPING OF THE SAHWAVE AND NIGHTINGALE RANGES, PERSHING COUNTY, NEVADA

Nicholas Van Buer Department of Geological and Environmental Sciences, Stanford University 450 Serra Mall, Bldg. 320, Stanford, CA 94305-2115

A version of Chapter 2, with Plates 1–8 recompiled into a single reduced-scale map, has been submitted to the Nevada Bureau of Mines and Geology for publication as an Open- File Report. NBMG does not claim copyright for its publications. 24

ABSTRACT

Geologic maps and cross sections at 1:24,000 scale were produced for eight 7.5´ quadrangles covering the Sahwave and Nightingale ranges in Pershing County, Nevada.

This mapping reveals the presence of Cretaceous magmatic arc rocks, including a large, concentrically zoned intrusive complex, which intrudes deformed Triassic/Jurassic metasedimentary rocks. The Mesozoic units are in turn intruded by Tertiary dikes and capped by a Tertiary unconformity overlain by Oligocene to Miocene volcanic and sedimentary strata. North-south normal faults cut all these rocks and are responsible for the uplift of the Sahwave and Nightingale ranges, which together form a large horst that is warped into a gentle syncline.

INTRODUCTION

The part of the northern Sierra Nevada Batholith that crosses the NW Basin and

Range is in many places obscured by Cenozoic cover, but relatively continuous outcrop of granitic rocks occurs in the area of the Sahwave and Nightingale Ranges, about an hour northeast of Reno (Fig. 2.1). Much of the previous mapping in this region was at reconnaissance scale (often 1:250,000) or focused on specific areas of mineralization or

Cenozoic and Quaternary history. To elucidate the structural and petrologic history of this part of the Sierra Nevada Batholith in greater detail, the Sahwave and Nightingale

Ranges were mapped at 1:24,000 scale, with an emphasis on portraying and defining basement structures. As a result of this mapping, this region was discovered to expose the only Late Cretaceous, large, zoned intrusive suite (similar to coeval intrusions in the eastern Sierra Nevada such as the Tuolumne Intrusive Suite) yet documented in the NW Nicholas J. Van Buer Figure 2.1.

123º OREGON 122º 121º 120º 119º 118º CALIFORNIA OREGON NEVADA 42º

KLAMATH

RANGE 41º Gerlach

Empire LovelockB A S I N 40º Nixon area of A N D Fig. 2.2 Reno R A N G E Lake Tahoe

S I E R R A N E V A D A 39º

Sacramento

G R E A T V A L L E Y

CALIFORNIANEVADA

38º

37º

Explanation of Map Units

Cenozoic cover

Mesozoic intrusions 36º Other Pre-Cenozoic rocks

Geologic map compiled mainly from Jennings et al., 1977; Irwin and Wooden, 2001. 50 km

Figure 2.1. Mesozoic intrusions of the Sierra Nevada Batholith (bounded by dotted lines) continue northwards into the Basin and Range, where they are partially obscured by Cenozoic cover. The study area has some of the most continuous outcrops available. 26

Basin and Range, providing a solid basis for depicting the Cretaceous batholith as trending NNE across the northwestern Basin and Range. This chapter primarily consists of eight 7.5´ quadrangle maps with cross sections (Plates 1–8), as prepared under the requirements of the U.S. Geological Survey EDMAP program, which provided partial support for this endeavor. The text of this chapter provides context for the maps and summarizes the geologic history of the region. The detailed petrology, geochemistry, and geochronology of the intrusive rocks are treated more thoroughly in Chapter 3 of this thesis and have been previously published (Van Buer and Miller, 2010).

GEOGRAPHY OF THE STUDY AREA

The mapped area covers most of the Sahwave and Nightingale mountain ranges, an area of almost 1200 km2 (Fig. 2.2). This area is about an hour northeast of Reno, and is accessible from the towns of Nixon, Gerlach/Empire, and Lovelock (Fig. 2.1). No part of the map area is more than half a day’s walk from a decent dirt road. The Sahwave and

Nightingale Ranges are about 50 km long and trend north-south (Fig. 2.2). Each range has a maximum relief of about 1 km, with elevations in the map area ranging from 1149 m at the lowest part of the Winnemucca basin to 2278 m at Juniper Mountain in the northern Sahwave Range (Figs. 2.2, 2.3). The land between the ranges, known as Sage

Hen Valley, lies 400–500 m higher than the basins to the east or west, and the topography of each range is generally steeper on the side facing away from this upland valley (Fig.

2.2). To the east and west, Granite Springs Valley and Winnemucca Valley are much broader than Sage Hen Valley and contain playas (Fig. 2.2), although Winnemucca

Valley held a more substantial lake as recently as 1945 (Zones, 1961). 27

Geologic K U M I V A maps consulted: Bluewing D Flat V A L L E Y

W M T S. B L U E W I N G

B J Juniper Pass 1 2

W S R S

S E L E N I T E R A N G E A E R T E N I S E L

40° 30’ N Juniper

Adobe Flat Mtn.

3 4 5

S A G E

H E N 6 7 8 V A L L E Y

Nightingale Mining

N I G H T District I N G A L E R A N G E S A H W A V E R A N G E 40° N

W I N N E M U C C A L A K E B E D 5 km

119° 30’ W 119°

119° W 119° G R A N I T E S P R I N G S V A L L E Y E L L A V S G N I R P S E T I N A R G L L A K E R A N G E T R U C K E E R A N G E Figure 2.2. Physical geography of the study area. The eight mapped 7.5’ quadrangles are 1 - Jayhawk Well, 2 - Juniper Pass, 3 - Tohakum Peak NE, 4 - Tunnel Spring, 5 - Bob Spring, 6 - Tohakum Peak SE, 7 - Sage Hen Spring, and 8 - Bluewing Spring. Geologic maps consulted: B - Bonham, 1969; D - Dasher, unpublished mapping; J - Jennings, 1977; R - Rai, 1969; S - Stager and Tingley, 1988; W - Whitehill, 2009. 28

Figure 2.3. View south towards Juniper Pass (center) through the northern Sahwave Range (high point Juniper Mountain) from the Bluewing Mountains, showing typical vegetation and outcrop patterns. Foreground outcrops are Triassic/Jurassic metasedimentary rocks, which are overlain unconformably by the Tertiary rhyolite flow which makes up the orange hills at left midground. Most of the Sahwave Range is Cretaceous granodiorite, which intrudes metasedimentary rock near the left edge of the photo, just above the highest orange hill. Adobe Flat, a playa, can be seen in the left background. Juniper Pass appears to be an antecedent stream gap. 29

The most rugged parts of the map area, in the western Nightingale Range and the northern Sahwave Range (Fig. 2.2), bear relatively good rock exposure, although slopes are often obscured by colluvium or scree (Fig. 2.3). Areas of lower relief, such as the southern Sahwave Range (Fig. 2.2), however, are often blanketed by a thick layer of grus, and, aside from a few rocky tors, dikes comprise a majority of the competent outcrop.

Vegetation is mainly sparse scrub, often dominated by sagebrush, and does little to obscure rock exposure (Fig. 2.3). A few juniper and mountain mahogany grow at the higher altitudes, but the only real trees are four cottonwoods at the mouths of two canyons in the northeastern Sahwave Range.

PREVIOUS WORK

The area was first described by Clarence King’s Fortieth Parallel Survey in 1867, but this was a brief reconnaissance only. They concluded that the Mesozoic metasedimentary rocks in this area were deposited unconformably over Archaean granitic basement (King, 1878). Tungsten was discovered in the Nightingale Mining District in

1917, leading to a number of local, economically-motivated studies (Fig. 2.2; Hess and

Larsen, 1922; Smith and Guild, 1942; East and Trengrove, 1950; Rai, 1969; Fanning,

1982; Stager and Tingley, 1988). Bonham (1969) and Johnson (1977) mapped Washoe and Pershing counties, respectively, at 1:250,000 scale, providing the first solid regional geologic map of the area, but did not distinguish separate plutonic units (Fig. 2.2). Other work in the region has focused on Oligocene paleovalleys (Faulds et al., 2005),

Quaternary lake deposits (e.g., Russell, 1885; Broecker and Orr, 1958; Zones, 1961;

Benson and Thompson, 1987), wave-cut cave archaeology and geology (e.g., Orr, 1956; 30

Sears and Roosma, 1961), and Holocene fault scarps (Wesnousky et al., 2005). The most recent mapping in the area was undertaken by Whitehill (2009), who investigated the

Tertiary volcanic and sedimentary strata as well as the extensional faults that have uplifted and warped the Sahwave and Nightingale ranges (Fig. 2.2). Other compilations containing geochemical data (but no geologic mapping) from basement rocks in the area include Zirkel (1876), Smith et al. (1971), Farmer and DePaolo (1983), and Wooden et al. (1999).

GEOLOGIC SETTING

The geologic history established in greater region surrounding the map area begins with Late Triassic (Norian) to earliest Jurassic deposition of deep marine sediments of the Auld Lang Syne Group (“basinal terrane” on Fig. 2.4; e.g., Burke and

Silberling, 1973; Johnson, 1977). Though generally unfossiliferous, these strata have been dated by the rare occurrence of ammonite fossils (Burke and Silberling, 1973).

These basinal strata are mostly fine-grained siliciclastic units with scattered lenses of coarser material and rare carbonate, and are essentially submarine fan deposits (Burke and Silberling, 1973; Speed, 1978). They are regionally metamorphosed to subgreenschist or lower greenschist facies, forming slates and phyllites, and are colloquially known as “the mudpile” due to their general lack of easily recognizable stratigraphy (Willden, 1964; Bonham, 1969; Johnson, 1977). The base of the unit is not exposed near the study area, where stratigraphic thicknesses over 6 km have been measured in the Triassic to Jurassic deposits (Burke and Silberling, 1973; Speed, 1978).

To the east, correlative strata are thinner, where they form the uppermost unit of a 31

OR Map explanation CA NV no data basinal terrane

Black Rock arc MSNI other early Mz arc

Black Rock Arc LFTB Triassic shelf deposits, overlapping Paleozoic basinal allochthons to the east Triassic MSNI important thrust faults study LFTB shelf area and terrane no data underlying early Mesozoic arc rocks Paleozoic

allochthons MSNI

LFTB

Figure 2.4. Terranes of the NW Basin and Range at the end of the Jurassic, before intrusion of the Cretaceous Sierra Nevada Batholith (modified from Martin et al., 2010). Only the frontal thrust of the Luning-Fencemaker Thrust Belt is shown (labeled LFTB on map); deformation associated with this belt extends throughout the basinal terrane (Oldow, 1984). MSNI = Mojave-Snow Lake-Nevada-Idaho Fault, active in the early to mid-Cretaceous (Wlyd and Wright, 2001, 2005). Terranes west of this fault not shown. 32 sequence of Triassic shelf deposits, which in turn overlie Permo-Triassic volcanic rocks and deformed Paleozoic deep marine strata of the Golconda and Roberts Mountains allochthons (Fig. 2.4; e.g., Silberling and Wallace, 1969; Speed, 1978). Upper Paleozoic and lower Mesozoic volcanic rocks northwest of the basinal terrane, referred to as the

Black Rock terrane, have a volcanic island arc affinity (e.g., Russel, 1984; Wyld, 1990).

These regional relationships suggests that the Auld Lang Syne basin may have formed in a back-arc extensional environment (Fig. 2.4; e.g., Quinn et al, 1997; Wyld, 2000).

The exact history of deformation recorded in these basinal sediments varies somewhat from place to place, but generally includes the formation of NE-SW-striking penetrative slaty foliation in the argillaceous units during Jurassic deformation (e.g.,

Oldow, 1984; Thole and Prihar, 1998; Wyld et al., 2001). Folding associated with the primary foliation is generally SE-vergent, as are thrust faults within a regional zone of deformation known as the Luning-Fencemaker Thrust Belt (Fig. 2.4; Oldow, 1984).

West of the basinal terrane, Wyld and Wright (2001) document an early to mid-

Cretaceous shear zone with major right-lateral strike slip (~400 km) and minor overprinting shortening (Fig. 2.4). This shear zone is inferred to be part of a regional fault system active at this time, the Mojave-Snow Lake-Nevada-Idaho (MSNI) shear zone

(Wyld and Wright, 2005), and projects just west of the study area (Fig. 2.4; Wyld and

Wright, 2001). Arc magmatism in the region was most voluminous from ca. 110 to 85

Ma (Fig. 2.1; e.g., Smith et al., 1971; Barton et al., 1988; Van Buer and Miller, 2010).

These mid- to late Cretaceous intrusions are associated with relatively narrow (up to kilometer scale) contact metamorphic aureoles, locally up to amphibolite hornfels, characterized by coeval deformation that overprints older fabrics (e.g., Compton, 1960; 33

Ciavarella and Wyld, 2008). Cretaceous deformation in the region is otherwise relatively minor, although there is local evidence for extensional faulting in the early

Cretaceous (e.g., Quinn et al., 1997; Wyld et al., 2001; Martin et al., 2010).

Magmatism was followed by late Cretaceous to early Tertiary erosion, and a widespread unconformity was developed across the region and overlain by Eocene to

Miocene volcanic and sedimentary strata (e.g., Stewart, 1998; Henry, 2008; Van Buer et al., 2009). In the NW Basin and Range, east-west extension began in the mid-Miocene, ca. 15-12 Ma, and continues to the present day, manifested by the formation of tilted, roughly north-south trending ranges, usually bound by a normal fault on one side

(sometimes both) and separated by deep Neogene basins (Fig. 2.1; e.g. Stewart, 1998;

Colgan et al., 2006; Lerch et al., 2008; Whitehill, 2009). Where extensional fault slip and associated tilting has been sufficiently great to rotate normal faults to shallow dips, a second (or even third) generation of steeper faults has developed (e.g., Proffett and

Dilles, 1984; Surpless et al., 2002). Though externally drained in the early Tertiary (e.g.,

Henry, 2008; Cassel et al., 2009), faulting-induced basin and range formation, in conjunction with a the rain shadow effect of the Sierra Nevada, has led most of the Basin and Range Province (the “Great Basin”) to become internally drained, with streams ending in playas, marshes, and occasionally, lakes. In the Pleistocene, the deeper basins of the NW Basin and Range were intermittently flooded by fresh water, including a regionally extensive lake system, Lake Lahontan. This lake filled to a maximum depth of

~280 m (shoreline elevation 1330 meters above sea level), last reached about 13,000 years ago (Benson and Thompson, 1987). This lake has left fine-grained lacustrine deposits, beach deposits, wave-cut terraces and caves, spits, perched deltas, and tufa 34 deposits (which often cement the previous formations into durable beachrock) below its highstand elevation (e.g., Russell, 1885; Benson, 2004).

METHODS

Mapping in the field was carried out using U.S. Geological Survey 7.5´ quadrangle topographic maps (1:24,000 scale) as base maps. Aerial photography

(U.S.G.S. 1 m resolution orthophotoquads) or satellite imagery (National Agricultural

Imagery Program 1 m resolution color orthorectified images), overlain with topographic map data, was consulted in the field while mapping, and again when digitizing in

ArcMap. For example, dike locations plotted in the field were often adjusted up to about

25 m (~1 mm on the map) upon digitization if their location in the digital image could be ascertained more precisely. (Note that image orthorectification is often only accurate to within ~6 m; Farm Service Agency, 2010.) Contacts and faults are generally plotted with solid lines only if their existence is certain and their position is thought to be accurate within 25 m; otherwise dashed lines are used (or dotted lines if the contact is concealed). In grus-blanketed areas, dashed contacts were sometimes mapped based on analysis of the soil, checked by occasionally digging small pits with my rock hammer.

Some of the contacts between Quaternary units, such as between sand dunes and alluvium, or between alluvium and lake deposits, were mapped primarily from digital imagery, though field checked in areas near rock outcrop—these contacts are never represented with solid lines. Previous geologic mapping (Fig. 2.2) was frequently consulted (especially Whitehill’s 2009 work emphasizing the Tertiary history) but contacts were never copied directly, and my interpretation sometimes differs from that of 35 previous authors. Some strikes and dips, however, were transferred from earlier maps, particularly in some parts of the Tertiary units (from Whitehill, 2009) and in parts of the

Blue Wing Mountains (Fig. 2.2; Sarah Dasher, unpublished mapping). My own structural measurements of obscure planar features, such as weak magmatic foliation, generally represent a field average of two or three measurements.

Because contact attitudes are rarely directly measurable in the area, attitudes shown in cross section are generally estimated from map patterns by structure contouring.

Where no other information is available, normal faults are assumed to have formed at 60° dip and tilted the same amount as the titled Tertiary strata they cut.

LITHOLOGIC UNITS

Basement rock units in the Sahwave and Nightingale ranges include early

Mesozoic metamorphic rocks and a number of Cretaceous plutonic rocks that intrude them. The plutonic rocks can be divided into two groups: rocks with an easily visible, roughly north-south, subvertical foliation (either magmatic or sub-solidus), and a set of younger, concentrically nested comagmatic intrusions with weak or contact-parallel magmatic foliation, referred to as the Sahwave Intrusive Suite. The Mesozoic rocks are beveled by an unconformity overlain by Oligocene and Miocene volcanic and sedimentary strata. Normal faulting, beginning in the Miocene, created the present structure of the Sahwave and Nightingale ranges and their flanking basins, which have been mantled in an array of Quaternary deposits. Map units will be discussed from oldest to youngest.

Triassic/Jurassic Metasedimentary Rocks

The oldest rocks in the map area, intruded by the Cretaceous Sahwave batholith, are mostly metamorphosed mudstone/shale with interbedded sandstone layers and lenses

(JTrm on Plates 1–8). A few discontinuous, 10–100 m thick, coarsely crystalline marble layers are present south of the batholith, but calcareous horizons are rare in the metamorphic rocks to the north. These rocks have been identified as submarine fan sediments belonging to the Triassic to lower Jurassic Auld Lang Syne Group (Johnson,

1977). A very limited number (n = 6) of restored flute-cast measurements indicate paleocurrents roughly from the southeast, which is consistent with derivation from the continental shelf that lay in that direction (Fig. 2.4). A particularly thick lens of quartzite, probably the remains of a submarine channel, holds up the peak of the Blue Wing

Mountains (Fig. 2.2).

Away from the Sahwave batholith, metamorphic grade is subgreenschist to lower greenschist facies, and original bedding can be clearly seen (though it is often folded).

Adjacent to the batholith, Triassic/Jurassic strata are metamorphosed to siliceous hornfels or biotite schists, and bedding is often tightly to isoclinally folded or transposed with a sub-vertical axial-planar foliation that is broadly parallel to the intrusive contact. A strong sub-vertical mineral lineation is also present within ~100m of the intrusive contact, although it is often obscured by a subparallel intersection lineation. With some exceptions, the generally quartzofeldspathic composition of the metasedimentary rocks is not conducive to the growth of diagnostic minerals besides white mica and occasional biotite. The Nightingale mining district (Fig. 2.2) contains a number of skarn deposits in the contact aureole of the batholith where calcareous layers have been metamorphosed, 37 yielding grossular/almandine, clinozoisite/epidote, and more rarely tremolite, wollastonite, diopside, and scheelite, in addition to the standard quartz, ±albite, and calcite. White mica pseudomorphs after andalusite (rod-shaped) and cordierite (mouse- turd-like), as well as cordierite-shaped voids, can be found in some of the more pelitic layers in the Bluewing Mountains (Fig. 2.2) near the northern margin of the batholith.

However, large (to 5 cm) andalusite crystals remain intact in at least one area about 1 km from the northern contact (Plate 2), growing in random orientations that cut across the foliation.

Early Intrusive Units

Power Line Intrusive Complex: The oldest intrusive unit, dated at 105 Ma (Van Buer and Miller, 2010; see chapter 3 for discussion of geochronology), is informally referred to as the Power Line intrusive complex (Kpl). Occupying the northwestern Nightingale

Range (Fig. 2.2), it is predominantly a medium-grained biotite hornblende granodiorite with 5–10 mm K-feldspar phenocrysts. This unit also includes many unmapped dikes and pods of darker granodiorite and diorite ranging from centimeters to hundreds of meters in dimension. Some of these are fine grained, weathering to a blue-grayish color, but all subunits share a similar, generally north-south oriented, solid-state foliation that tends to be more pronounced towards the eastern side of the Nightingale Range (Plates 3,

6). Although many of the finer-grained mafic enclaves appear to demonstrate magma mingling, relationships among these subunits are somewhat obscured by poor outcrop and the solid-state foliation. In thin section, the foliation is defined by biotite strung out along wavy foliation planes, and the sense of shear, if any, is unclear, as the rock bears no 38 discernible lineation. Biotite and quartz appear to have been largely recrystallized, but feldspars remain intact, displaying distinct undulatory extinction, suggesting solid-state deformation at temperatures of ca. 400–450 °C or hotter, depending on strain rate

(Passchier and Trouw, 2005). This unit also contains many large inclusions or apophyses of metamorphic rock, mostly 5–200 m in length but including a 4-km-long possible roof pendant (Plates 3, 4). These are generally elongated in map view, and foliated steeply, concordant with the foliation of the Power Line complex (Plates 3, 4).

Selenite Granodiorite: In the very northwestern corner of the study area and throughout the southern Selenite range (Fig. 2.2) is a distinct granodiorite here referred to as the

Selenite Granodiorite (Kse) after the “Selenite pluton” of Smith et al. (1971). This unit, dated at ca. 96 Ma (Van Buer and Miller, 2010), has a conspicuous, generally north-south magmatic foliation defined by the alignment of euhedral plagioclase and hornblende phenocrysts in rock with a hypidiomorphic igneous texture. Polysynthetic twinning in the plagioclase is frequently visible to the unaided eye.

Sahwave Intrusive Suite

The metamorphic rocks, the Power Line complex, and the Selenite Granodiorite are intruded by the large (~1000 km2), zoned Sahwave Intrusive Suite, which comprises most of the basement outcrop in the map area. This intrusive suite consists of three concentric intrusive units and a distinct lobe-forming unit that stretches across the central

Nightingale Range (Fig. 2.2).

Granodiorite of Juniper Pass: The outermost and oldest intrusive unit is a medium-to- coarse-grained equigranular biotite hornblende granodiorite referred to as the

Granodiorite of Juniper Pass (Kjp) due to its excellent exposure at Juniper Pass in the northern Sahwave Range (Fig. 2.2, 2.3). Dated at ca. 93 Ma (Van Buer and Miller,

2010), the Granodiorite of Juniper Pass is identified by its conspicuous 4–8 mm euhedral biotite crystals. Additionally, large hornblende phenocrysts are common around the periphery of this intrusion, giving the rock a characteristic “dalmatian” appearance.

Hornblende and sphene are both present throughout the Sahwave batholith, but only in the Granodiorite of Juniper Pass does the hornblende form crystals notably larger than the

1–3 mm euhedral sphene. In detail, the mineral proportions and color index of this unit varies quite a bit; in places it can be classified as a tonalite or a quartz diorite.

Gradational compositional variation can sometimes be seen across large outcrops; more rarely, internal contacts can be discerned where slightly lighter and darker phases occur together. In a few places, straight or wavy compositional layers 1 cm – 1 m thick are bounded by sharp contacts. Many of these internal structures are subtle, and only readily seen in fresh outcrop, so it is possible that they are fairly pervasive. Mafic enclaves are found throughout the unit, but are only common within 1–2 km of the exterior contacts.

Enclaves are typically 5-30 cm in length and flattened by a ratio of 2:1 to 5:1 or more.

Mafic schlieren are common in the same region. The Granodiorite of Juniper Pass has a discernible magmatic foliation that is defined by the alignment of mafic minerals and sometimes subhedral plagioclase, which is generally similar to the alignments of mafic schlieren and mafic enclaves as well. Magmatic foliation tends to be strongest near the outer contact, which it often parallels. 40

Where contact orientation is evident, it tends to be steeply dipping and subparallel to magmatic foliation, but there are a couple of notable exceptions. These are the contacts along the two largest metamorphic blocks or pendants at the southern end of the

Sahwave Range (Plates 7, 8), and the shallow contact where the Granodiorite of Juniper

Pass underlies the Power Line Complex in the northwestern Nightingale Range (Plates 3,

4). It is not clear, however, whether these metamorphic pendants represent the true roof of the intrusion. In the northwestern Nightingale range, the low-angle portion of the contact terminates westward as the top contact of a horizontal dike of the Granodiorite of

Juniper Pass intruded across the foliation of the Power Line complex (Plate 3), suggesting that the contact in this area may simply surround a flap of wall rock that was in the process of being stoped.

Granodiorite of Bob Spring: The Granodiorite of Juniper Pass grades inward to the more felsic and uniform Granodiorite of Bob Spring (Kbs), dated at ca. 93 Ma (Van Buer and Miller, 2010), a medium-grained biotite granodiorite or granite characterized by seriate K-feldspar phenocrysts up to ~ 2 cm. In the field, this gradational contact is arbitrarily mapped where large K-feldspar phenocrysts become more conspicuous than large biotite crystals. Biotite in the Granodiorite of Bob Spring is more homogeneously distributed, and generally no larger than 1 mm. The K-feldspar phenocrysts are poikilitic, mostly surrounding plagioclase and biotite, and are occasionally sieve textured and difficult to see. In general, Kbs is finer grained towards its center, and K-feldspar phenocrysts are less common. The Granodiorite of Bob Spring bears equant quartz grains that are generally only about 1 mm in size but reach 3–5 mm in the southern part. 41

In the southern Sahwave Range this unit crops out in a few rocky tors, which in some cases appear to have gained their greater resistance to weathering by hydrothermal silicification. Mafic minerals are often badly chloritized and feldspars show signs of sericitization. Foliation in this unit is usually absent or at least too weakly defined to measure.

Sahwave Granodiorite: The Sahwave Granodiorite (Ks), a K-feldspar-megacrystic biotite granodiorite dated at ca. 89 Ma (Van Buer and Miller, 2010), intrudes the central part of the Granodiorite of Bob Spring along a generally shallowly dipping contact that is sharp on the north side but gradational along its south side (Plates 5, 8). K-feldspar megacrysts are 2-4 cm across, moderately poikilitic, and more abundant (usually 1–5% by volume) than in the Bob Spring Granodiorite. The abundance of K-feldspar megacrysts can vary greatly from place to place and at outcrop scale it is not uncommon to see distinct stringers and pods enriched in K-feldspar megacrysts, rarely up to as much as ~20%. These can vary from centimeters to meters in scale, do not have particularly sharp contacts, and are generally surrounded by rock that is more equigranular and leucocratic than average, suggesting that K-feldspar crystals may have been concentrated from the magma by flow segregation, or perhaps that leucocratic melt was filter-pressed out of fluid-softened areas of the magma. The Sahwave Granodiorite forms relatively bold outcrops compared to adjacent parts of the Granodiorite of Bob Spring, but the rock is uniformly crumbly and often spheroidally weathered.

School Bus Granodiorite: The Nightingale Range contains a distinct lobate unit referred to as the School Bus Granodiorite (Ksb), after outcrops in a canyon containing two defunct school buses. This unit, dated at ca. 91 Ma (Van Buer and Miller, 2010), is a relatively leucocratic granodiorite, distinguished by scattered 1–2 cm K-feldspar phenocrysts and 3–6 mm biotite flakes. Unlike the main part of the Sahwave batholith, this lobe does not appear to be any more mafic around its outer edge, and is, in fact, remarkably homogeneous. Magmatic foliation is not generally distinguishable. The

School Bus Granodiorite intrudes both the Power Line complex and the Granodiorite of

Juniper Pass along sharp, vertical contacts that are fairly irregular at the map scale (Plates

3, 4, 6, 7). Where it intrudes the Power Line complex, the units are often separated by metamorphic screens and blobs 20-200 m thick (Plates 3, 4).

Minor Cretaceous Intrusives

The southern part of the Granodiorite of Juniper Pass contains a number of diorite bodies (Kd), varying from tens of meters to over a kilometer in scale (Plate 7). These fine- to medium-grained intrusions frequently contain ~5 mm euhedral plagioclase phenocrysts, and sometimes acicular hornblende crystals as well. These diorite bodies appear to be coeval with the Sahwave batholith, often showing magma mingling and mixing structures such as lobate and interfingering contacts, streaky fine-scale intermingling, and outcrop-scale continuous compositional variation indicative of wholesale mixing.

Additionally, the Sahwave batholith and its country rocks are pervasively intruded by a series of leucocratic dikes and sills (Kap) that tend to be more resistant to weathering 43 than the surrounding country rocks. Most of these dikes demonstrate wide variations in grain size between aplite and pegmatite textures, often showing evidence for repeated intrusion. The pegmatites are generally muscovite bearing, and may also contain tourmaline (schorl) and rarely garnet. The dikes range from 1 cm to 100 m in thickness and generally strike north-south, but in the Sahwave Range vary systematically between

NNW trending in the south and NNE trending in the north. Dips are often moderately shallow, but only locally consistent in direction of dip. The most notable pegmatites, containing feldspar crystals up to 1 m in dimension, occur in the northeastern Sahwave

Range, where they define a domal structure (Plate 5). Variations in the orientation of these dikes are hypothesized to be due to perturbations in the principal stresses as a result of thermal contraction of the batholith. Another notable concentration of leucocratic dikes exists in the Nightingale Range, intruding the Power Line complex (Plate 3). These dikes crosscut the solid-state foliation of the Power Line complex and are occasionally composite, containing a phase with scattered large K-feldspar and biotite phenocrysts, suggesting that they may be genetically related to the School Bus Granodiorite, which could underlie this area beneath the current level of exposure (Plate 3). Pegmatite dikes cutting the metamorphic rocks along the margins of the batholith are frequently folded and boudinaged in the foliation, which is subparallel to the country-rock contact. A few broader leucogranite intrusions (Klg), which are more uniform in grain size and contain minor biotite, are present near the southern margin of the batholith (Plates 6–8).

Tertiary Dikes

The Sahwave batholith is cut by a few generations of younger dikes, ranging from rhyolitic to basaltic in composition, which tend to be more resistant to erosion than the surrounding rock. Although these dikes have not been dated, they cut all aplite dikes.

Their generally fine-grained nature suggests that they are substantially post-magmatic, and they are compositionally similar to volcanic rock units in the overlying Tertiary section. In order of intrusion, (as determined by cross-cutting relationships), the dike units are as follows: (1) Tb3 – dark, greenish-gray, aphanitic, aphyric basalt dikes. (2)

Td2 – very fine-grained gray diabase dikes. (3) Tb2 – pyroxene ± olivine phyric basalt dikes (phenocrysts < 1 mm; often altered). (4) Tb – plagioclase phyric basalt dikes. The plagioclase phenocrysts are often of the size and appearance of rolled oats (~ 4 mm). (5)

Tf – rhyolite dikes, generally flow-banded and a minimum of 10 m thick. Often they contain large quartz and biotite crystals that may be xenocrysts scavenged from the surrounding granodiorite. In areas with shallow inferred Tertiary paleodepth, rhyolite intrusions take the form of equant plugs 300–400 m across (Plates 1, 4). (6) Td – coarser diabase dikes, containing plagioclase phenocrysts up to 3 mm, also some hornblende in the matrix. Though dark grey when fresh, these dikes weather to a distinctive spearmint- green hue. Unlike all the previous dikes, which generally trend WNW-ESE, these dikes trend NNE-SSW, paralleling the range-bounding normal faults (Plate 5). These dikes dip steeply towards the range-front fault, and have likely been back-tilted from an initially vertical orientation (Plate 5). They are the only dikes whose orientation may bear a kinematic relationship to east-west Basin and Range extension—the WNW-ESE orientation of the earlier dikes either reflects a pre-extensional stress orientation or the 45 influence of the weak but approximately parallel magmatic fabric of the Cretaceous intrusive suite the dikes intrude (e.g., Plate 2).

Tertiary Strata

All metamorphic and plutonic rocks are unconformably overlain by Oligocene and Miocene volcanic and sedimentary rocks, which have been fairly throroughly described by Faulds et al. (2005) and Whitehill (2009). The following description includes my own observations as well. The volcanic units range from basalt flows to silicic ignimbrites, and the interbedded sediments (not thoroughly lithified) range from landslide deposits and fanglomerates to lacustrine clays (cf. Whitehill, 2009). Other than an occasional boulder lag or conglomerate, the basal Tertiary unconformity is frequently directly overlain by volcanic or volcaniclastic rocks. In the southern Nightingale Range

(Fig. 2.2), the basal volcanic rocks, which fill a paleochannel, are Oligocene ignimbrites dated at 31–25 Ma (sanidine 40Ar/39Ar; Faulds et al., 2005), but elsewhere in the map area they are Miocene basalts dated at 18–14 Ma (whole-rock 40Ar/39Ar; Whitehill, 2009), volcaniclastic units, and locally rhyolite flows. Just northwest of the map area, however, lacustrine deposits form the base of the section (Whitehill, 2009). The basal strata appear to predate normal faulting in the area, but younger volcanic rocks, dated at 14–12 Ma

(whole-rock 40Ar/39Ar; Whitehill, 2009), and volcaniclastic sediments may form thicker sequences at the base of nascent fault-bound extensional basins (Whitehill, 2009). In the southwestern Sahwave Range, primarily non-volcaniclastic sedimentary rock, including fanglomerate, sandstone, and shale (ca. 12–9 Ma) has been exposed by later faulting 46

(Plate 8; Whitehill, 2009), and likely represents the sort of material which has continued to fill the adjacent Granite Springs and Winnemucca Cenozoic basins until the present day (Fig. 2.2). This mapping distinguished only three Tertiary units: (Tv) primarily volcanic and volcaniclastic rock; (Ts) primarily non-volcaniclastic sedimentary strata, ranging from alluvial to lacustrine deposits, tilted, but often not fully lithified; and (Tc) a distinctive (though poorly exposed) boulder conglomerate in the east-central Nightingale

Range (Plate 4) that is presumed to be Tertiary because it contains metamorphic and volcanic clasts from units not present in the current watershed.

Quaternary Units

About half of the map area is covered by Quaternary deposits. Areas of basement rock covered with a thin veneer of alluvium, colluvium, or scree, were mapped according to the basement lithology if it could be reasonably inferred from soil composition or exposure in frequent gullies. Quaternary deposits were only mapped where they appeared to be more than about 2 m thick and 25 m wide. Older alluvial terraces are often preserved, but for the puposes of this study all alluvial fan and stream deposits were mapped as a single unit (Qa). Slumps, talus slopes, and other mass wasting deposits

(Qls) are common in the steeper parts of the map area.

Winnemucca Valley and Kumiva Valley (Fig. 2.2) both contain Pleistocene lake deposits (Ql). Extensive deposits related to ancient Lake Lahontan occur throughout

Winnemucca Valley at elevations below about 1340 m (Fig. 2.2; Plates 3, 6), including coarse sand and gravel beach deposits, delta deposits, and fine-grained lake deposits

(mapped as a separate unit, Qp, where reworked in the modern playa). Alluvial deposits 47 surficially terraced by wave action are nearly indistinguishable from thicker beach deposits, and are also mapped as Ql. Tufa is deposited extensively as a cementing agent, as a veneer on rock surfaces, and as hummocks or towers up to ~20 m high (Plates 3, 6).

Tufa-cemented clastic rocks are generally mapped as Ql, whereas freestanding tufa deposits are mapped as a separate unit, Qt. Lacustrine deposits at 1340 m (e.g., in

Stonehouse Canyon, Plate 3) would appear to be about 10 m above the most recent 1330 m highstand level for Lake Lahontan (Benson and Thompson, 1987). This deviation is within the range reported throughout the Lahontan basin due to isostatic rebound and recent tectonic uplift; alternatively, these deposits could date from an earlier, higher highstand (e.g., Reheis et al., 2002).

In Kumiva Valley, the only significant lake deposit is a low, kilometer-long gravel bar about a kilometer south of Bluewing Flat (Fig. 2.2), which lies about 10 m above the elevation of the playa surface. This gravel bar is inferred to mark the ~1352 m highstand level of Lake Kumiva (cf. Mifflin and Wheat, 1979). About 20 m higher absolute elevation than the highstand of Lake Lahontan (Benson and Thompson, 1987), this would have been a small isolated lake on the large island of a dozen or so ranges surrounded by Lake Lahontan (e.g., Russell, 1885).

Additionally, two active playas (Qp) were mapped. Bluewing Flat is very smooth, flat, and usually bone dry, but Winnemucca Lake received more substantial quantities of water in the recent past and consequently has a bit more topography—many areas might be better classified just as fine-grained lake deposits rather than an active playa surface. The water table seems to be very shallow, and the lowest parts of the playa can remain muddy (sometimes under a dry surface crust) quite late in the year 48

(don’t drive on this playa).

Eolian deposits (Qe) are among the youngest features in the area. This unit includes inactive, vegetation-stabilized dunes and sand sheets, mostly in Kumiva Valley

(Plates 1, 2), as well as active dunes, mostly in Winnemucca Valley (Plates 3, 6). Field checking indicates that some of the eolian features inferred from satellite imagery, especially in Kumiva Valley, are in fact rather thin veneers of wind-blown sand.

However, all such features have been retained for consistency, due to the practical difficulty of applying a robust minimum thickness limit. There are also plenty of dunes on the surface of Winnemucca playa (Fig. 2.2), which were not mapped due to low visual contrast against the playa sediments in satellite imagery.

STRUCTURAL DEVELOPMENT

Jurassic to mid-Cretaceous Deformation

The earliest record of deformation in the map area is preserved in fabrics developed in the Upper Triassic to lowest Jurassic metasedimentary rocks (JTrm) in the

Bluewing Mountains (Fig. 2.2). Here, foliation and the axes of outcrop-scale, tight to isoclinal folds trend NE-SW and dip steeply (Plate 2). Additionally, at map scale, the primary foliation of the western Blue Wing Mountains is folded into an open, recumbent, overturned to the SE syncline with parasitic folds in its core (Plate 2). These features are consistent with the SE-vergent style of Jurassic deformation in nearby parts of the

Luning-Fencemaker thrust belt (Fig. 2.4; Oldow, 1984; Thole and Prihar, 1998).

In contrast, the 105 Ma Power Line intrusive complex and the many metamorphic pendants it contains are marked by a strong, roughly north-south, subvertical solid-state 49 foliation. The 96 Ma Selenite Granodiorite shares a magmatic foliation of the same orientation. Without a clear lineation, these foliations remain kinematically ambiguous.

However, this deformation could be related to the early- to mid-Cretaceous MSNI shear zone of Wyld and Wright (2001, 2005), which shares approximately the correct orientation, location, and timing.

Emplacement of the Sahwave Intrusive Suite

These early units are all truncated by the Sahwave Intrusive Suite. Although foliation in the Powerline intrusive complex is sometimes locally perturbed near the

Sahwave Intrusive Suite contact, the most obvious deformation spatially associated with the Sahwave batholith appears in the adjacent metasedimentary rocks. These are locally foliated parallel to the steep intrusive contact with a strong subvertical lineation, suggesting that the wall rocks were flattened in pure shear and flowed ductilely downwards to accommodate the laterally expanding pluton. Boudinaged pegmatites of approximately the same age as the batholith (Van Buer and Miller, 2010) clearly demonstrate that ductile deformation was taking place in the contact aureole of the intrusion. In the Blue Wing Mountains (Fig. 2.2), along the northern edge of the batholith, the zone of contact-parallel foliation is only a few hundred meters wide (Plate

2), whereas to the southwest, in the Nightingale mining district (Fig. 2.2), foliation is subparallel to the contact over the entire region of exposed metamorphic rock, up to 5 km away from the intrusion (Plates 6, 7). This NW-SE foliation is anomalous compared to

NE-SW Jurassic structural trends in surrounding areas (Oldow, 1984). Re-intrusion of the School Bus Granodiorite so close to the contact already heated once by the 50

Granodiorite of Juniper Pass may be responsible for the more extensive ductile deformation (and mineralization) along the southwestern edge of the intrusion (Plates 6,

7), whereas the active intrusive contact retreated significantly from the northern margin over time (Plates 2, 5).

The steep contacts and magmatic foliation of the Sahwave Intrusive Suite suggest that it is not a sill-like intrusion. On the other hand, the presence of andalusite in its aureole limits its depth to less than ~ 10 km. Given its 40 km diameter at this depth, it must have been relatively flat-topped (constrained by the proximity of the surface). The downward extent of the batholith is not well defined by existing seismic or gravity data, but comparison to the oblique crustal arc sections exposed in southern California suggests that batholithic rocks such as these may extend to the base of the crust, although large distinct intrusions in the upper crust may overlie a complex zone of smaller, sheeted intrusions in the lower crust (Saleeby, 2003; Barth et al., 2008; Saleeby et al., 2008).

Conversely, seismic data from further north along the arc (Lerch et al., 2007) indicate low velocities compatible with tonalitic/granitic rocks down to only ~15 km, so the magmatic arc in NW Nevada may not encompass the entire crust, or may be more mafic at depth.

Even if the batholithic rocks studied here extend only to an additional 15 km depth, this intrusive suite still represents well over 10,000 km3 of batholithic volume.

Some fraction of this space may have been generated by ductile shouldering-aside and downward return-flow of the wall rocks, but this is frequently considered to be a relatively minor material transfer process in the upper crust (e.g., Paterson and Vernon,

1995). Major upper-crustal assimilation is ruled out isotopically (Van Buer and Miller, 51

2010). External contacts frequently dike into the metamorphic rocks and apparently surround stoped blocks, suggesting that stoping is at least a locally important process. If the kilometer-scale metamorphic outcrops in the southern part of the Granodiorite of

Juniper Pass (Plates 7, 8) represent stoped blocks rather than roof pendants, this would indicate a greater role of stoping in the intrusive process. Gradual tectonic opening could make some additional accommodation space. For example, right-lateral motion across the North American subduction zone in the Late Cretaceous (e.g. Engebretson et al.,

1985; Müller et al., 2008) could have produced significant localized extension if partitioned onto an intra-arc shear zone organized into right-stepping en-echelon segments (Tikoff and de Saint Blanquat, 1997; Tikoff and Teyssier, 1992). Such a shear zone has been identified in the Sierra Nevada (e.g., Busby-Spera and Saleeby, 1990;

Greene and Schweikert, 1995), but not yet in the study area.

Given the vast size of the Sahwave Intrusive Suite and the variations in lithology between and among its members, it seems likely to have been emplaced as a series of many intrusive pulses over its approximately four-million-year history (Van Buer and

Miller, 2010). However, smooth compositional variation and concentric arrangement within the gradationally zoned Sahwave Intrusive Suite suggests that individual batches of magma generally stayed hot long enough to partially mix with their successors. The

Granodiorite of Bob Spring, the School Bus Granodiorite, and the Sahwave Granodiorite are particularly homogeneous, suggesting that each may represent a single phase of rapid magma input into a large, partially molten magma chamber. The concentric arrangement of successively more homogeneous (and generally more differentiated) units also suggests that the system was warming over time, allowing larger and longer-lived magma 52 chambers to be formed at both the level of exposure and perhaps at the deeper level of magma production.

Mid-Miocene to Present Normal Faulting

Major normal faults bound the Nightingale and Sahwave ranges on the west and east sides, dipping west and east, respectively, such that the two ranges together form a horst block (cf. Whitehill, 2009). Still active today, the main faults exert strong control on topography, often truncating ridges and forming triangular facets, most spectacularly preserved in the southern Nightingale Range (Fig. 2.2, just left of the “N” in “Nightingale

Range”). These faults have back-tilted the strata in both their hanging walls and their footwalls, warping the rocks between the two ranges into a broad syncline, with minor faults in the hinge area (cf. Whitehill, 2009). In the central and northern Nightingale

Range, slip is actually partitioned between several, subparallel, smaller-offset faults

(Plates 3, 4), causing the range to bifurcate at its northern end (Fig. 2.2) where the faults take a more northeasterly trend (cf. Whitehill, 2009). This system of faults continues to the northeast of the Nightingale Range across Kumiva Valley, where the faults are largely concealed, but raise a few subtle hills (Fig. 2.2) and faint Quaternary scarps (Plates 1, 2).

The NE-striking, NW-dipping fault that uplifts the Bluewing Mountains (Fig. 2.2; Plate

2) appears to be a continuation of this system as well. In cross section (Plates 3, 4, 6), cumulative slip across the range-bounding faults varies from about 5 km in the south to about 3 km further north.

Faulting in the Sahwave Range appears to be separated into two kinematically distinct segments by an E-W accommodation fault (Plate 5). South of this fault there 53 appear to be two generations of N-S–trending faults (Plate 8). The first fault zone dips at a low angle (measured at ~33°; Whitehill, 2009), and is probably inactive at the surface today. It exhibits a fault-line scarp against a package of progressively tilted, ca. 12–9 Ma sediments in its hanging wall (Plate 8; Whitehill, 2009), but this may be simply related to the greater erosional resistance of granodiorite compared to partially lithified sediment.

The second fault system breaks ground east of the first fault (Plate 8), offsetting

Holocene alluvial terraces by as much as 3–4 m across an obvious fault scarp, and appears to dip more steeply (Whitehill, 2009; Wesnousky, 2005). To the south, this fault zone steps away from the range in a series of four left-stepping, en-echelon faults (Plate

8). These younger faults are responsible for exhuming the basin sediments exposed in the hanging wall of the older normal fault (Plate 8; Whitehill, 2009). These sediments are back-tilted ~30° (Plate 8), consistent with the older normal fault having originally formed at ~60° (Whitehill, 2009). Assuming a basin depth of ~2.5 km, as modeled from gravity data by Saltus and Jachens (1995), my cross sections show that the southern

Sahwave normal faults have a cumulative slip of about 9 km, with about a quarter of the total slip taken place along the younger faults (Plates 5, 8). Given that the southern

Nightingale Range has much more rugged topography than the Sahwave Range (Fig. 2.2) despite an apparently much smaller total slip, I speculate that the titling and uplift of the

Nightingale Range is, on average, younger than that of the southern Sahwave Range.

The two sides of the intervening syncline, which exhibit somewhat different styles of faulting, may not have formed entirely at the same time.

North of the accommodation fault, the Sahwave Range bends NNE (Fig. 2.2), and is bounded by normal faults on both sides. The eastern range-bounding fault exhibits a 54 clear Holocene scarp (Wesnousky, 2005), but there is no evidence for an older, low-angle fault. The most abrupt topography in the range switches to the west side instead of the east side (Fig. 2.2). Near its northern end, the range becomes narrower and lower, and is breached by what appears to be an antecedent stream gap, Juniper Pass (Fig. 2.3).

Gravity data also suggest that the northern part of Granite Springs Valley (Fig. 2.2) is shallower (~1.5 km) than southern part (Saltus and Jachens, 1995). I infer from these lines of evidence that the northern Sahwave Range is neither as tilted, nor as deeply exhumed on the east side, as the southern Sahwave Range. The Tertiary section has been completely eroded from the northern Sahwave Range, but if it has been uplifted as much as the basin has sunk, fault slip on the east side would only be ~4–5 km, and maybe half as much on the west side (cross sections on Plates 1, 2, 4, 5). From my cross sections

(Plates 2–8), I estimate the total extension across the Sahwave and Nightingale ranges north of the accommodation fault to be ~25–30%, and ~40–45% to the south of it.

SUMMARY

The geologic history of the Sahwave and Nightingale ranges can be summarized as follows:

1. Late Triassic to earliest Jurassic deposition of submarine fan sediments in a deep

water basin.

2. Late Jurassic southeast-vergent folding, low-grade metamorphism, and foliation

development related to the NE–SW-trending Luning-Fencemaker thrust belt.

3. Mid-Cretaceous intrusion of the Powerline intrusive complex and Selenite

Granodiorite during north-south fabric development (MSNI shear zone???). 55

4. Emplacement of the Sahwave Intrusive Suite (~93–89 Ma); generation of the

surrounding metamorphic and deformational aureole in the metasedimentary

rocks.

5. Erosion and exposure of the batholith.

6. Deposition of Oligocene to Lower Miocene volcanic and sedimentary strata

unconformably above Mesozoic basement. Intrusion of Tertiary dikes (?).

7. Mid-Miocene to present normal faulting and synclinal warping to generate the present

Sahwave and Nightingale Ranges. Additional Miocene volcanism. Continuing

erosion from the ranges and deposition in the basins.

8. Intermittent inundation of parts of Winnemucca and Kumiva valleys by Pleistocene

lakes and generation of present surficial deposits.

ACKNOWLEDGEMENTS

Partial support was provided by the U.S. Geological Survey National Cooperative

Mapping Program (EDMAP), NSF Tectonics grant 0809226, two Stanford McGee

Grants, and a GSA Student Research Grant. Thanks to my advisor, Elizabeth Miller, and my field assistants, Spencer Craven (who helped work out part of the dikes’ story),

Laainam “Best” Chaipornkaew, Silas Stafford, and Sarah Dasher (who helped map the

Bluewing Mountains). Thanks also for advice from several people at the Nevada Bureau of Mines and Geology and from Carrie Whitehill.

REFERENCES

Barth, A.P., Anderson, J.L., Jacobsen, C.E., Paterson, S.R., Wooden, J.L., 2008, Magmatism and tectonics in a tilted crustal section through a continental arc, 56

eastern Transverse Ranges and southern Mojave Desert: Geological Society of America Field Guide, v. 11, p. 101-117.

Barton, M.D., Battles, D.A., Debout, C.E., Capo, R.C., Christensen, J.N., Davis, S.R., Hanson, R.B., Michelson, C.J., Trim, H.G., 1988, Mesozoic contact metamorphism in the western United States, in Ernst, W.G., ed., Metamorphism and crustal evolution of the western United States, Rubey Volume 7: Englewood Cliffs, New Jersey, Prentice Hall, p. 110-178.

Benson, L.V., 2004, The Tufas of Pyramid Lake: U.S. Geological Survey Circular 1267, 14 pp.

Benson, L.V., and Thompson, R.S., 1987, Lake-level variation in the Lahontan Basin for the past 50,000 years: Quaternary Research, v. 28, p. 69-85.

Bonham, H.F., 1969, Geology and Mineral Deposits of Washoe and Storey Counties, Nevada: Nevada Bureau of Mines Bulletin, v. 70, 140 pp.

Burke, D.B., and Silberling, N.J., 1973, The Auld Lang Syne Group, of Late Triassic and Jurassic(?) age, north-central Nevada: Contributions to Stratigraphy, United States Geological Survey Bulletin 1394-E.

Broecker, W.S., and Orr, P.C., 1958, Radiocarbon chronology of Lake Lahontan and Lake Bonneville: Geological Society of America Bulletin, v. 69, p. 1009-1032.

Busby-Spera, C., and Saleeby, J.B., 1990, Intra-arc strike-slip fault exposed at batholithic levels in the southern Sierra Nevada, California: Geology, v. 18, p. 255-259.

Cassel, E.J., Calvert, A.T., and Graham, S.A., 2009, Age, geochemical composition, and distribution of Oligocene ignimbrites in the northern Sierra Nevada, California; implications for landscape morphology, elevation, and drainage divide geography of the Nevadaplano: International Geology Review, v. 51, p. 723-742.

Ciavarella, V., and Wyld, S.J., 2008, Wall rocks as recorders of multiple emplacement mechanisms—Examples from Cretaceous intrusions of northwest Nevada, in Wright, J.E., and Shervais, J.W., eds., Ophiolites, Arcs, and Batholiths: Geological Society of America Special Paper 438, p. 517-550.

Colgan, J.P., Dumitru, T.A., Reiners, P.W., Wooden, J.L., Miller, E.L., 2006, Cenozoic tectonic evolution of the Basin and Range Province in northwestern Nevada: American Journal of Science, v. 306, p. 616-654.

Compton, R.R., 1960, Contact metamorphism in Santa Rosa Range, Nevada: Geological Society of America Bulletin, v. 71, p. 1383-1416. 57

East, J.H., and Trengrove, R.R., 1950, Investigation of the Nightingale tungsten deposit, Pershing County, Nevada: U.S. Bureau of Mines Report of Investigations 4678.

Engebretson, D.C., Cox, A., Gordon R.G., 1985, Relative motions between oceanic and continental plates in the Pacific Basin: Geological Society of America Special Paper 206.

Fanning, D.J., 1982, Metamorphism and tungsten mineralization in the Nightingale Range, Pershing County, Nevada: [thesis; U Nevada Reno].

Farm Service Agency, 2010, National Agricultural Imagery Program (NAIP) Imagery: http://www.fsa.usda.gov/FSA/apfoapp?area=home&subject=prog&topic=nai accessed Feb. 29, 2011.

Farmer, G.L., DePaolo, D.J., 1983, Origin of Mesozoic and Tertiary granite in the western United States and implications for pre-Mesozoic structure; 1, Nd and Sr isotopic studies in the geocline of the northern Great Basin: Journal of Geophysical Research, v. 88, p. 3379-3401.

Faulds, J.E., Henry, C.D., Hinz, N.H., 2005, Kinematics of the northern Walker Lane; an incipient transform fault along the Pacific-North American Plate boundary: Geology, v. 33, p. 505-508.

Greene, D.C., Schweikert, R.A., 1995, The Gem Lake shear zone: Cretaceous dextral transpression in the Northern Ritter Range pendant, eastern Sierra Nevada, California: Tectonics, v. 14, p. 945-961.

Hess, F.L., and Larsen, E.S., 1922, Contact-metamorphic tungsten deposits of the United States: U.S. Geological Survey Bulletin 725, p. 245–309.

Irwin, W.P., Wooden, J.L., 2001, Map showing plutons and accreted terranes of the Sierra Nevada, California, with a tabulation of U/Pb isotopic ages: U. S. Geological Survey, Open File Report 01-0229.

Jennings, C., R. Strand, and T. Rogers, 1977, Geologic map of California, scale 1:750,000, Calif. Div. of Mines and Geol., Sacramento.

Johnson, M.J., 1977, Geology and Mineral Deposits of Pershing County, Nevada: Nevada Bureau of Mines and Geology Bulletin 89, 115 pp.

Lerch, D.W., Miller, E.L., McWilliams, M., and Colgan, J., 2008, Tectonic and magmatic evolution of the northwestern Basin and Range and its transition to unextended volcanic plateaus; Black Rock Range, Nevada: Geological Society of America Bulletin, v. 120, p. 300-311.

Martin, A.J., Wyld, S.J., Wright, J.E., and Bradford, J.H., 2010, The Lower Cretaceous King Lear Formation, northwest Nevada—Implications for Mesozoic orogenesis in the western U.S. Cordillera: Geological Society of America Bulletin, v. 122, p. 537-562.

Mifflin, M.D., and Wheat, M.M., 1979, Pluvial lakes and estimated pluvial climates of Nevada: Nevada Bureau of Mines and Geology Bulletin 94, 57 pp.

Müller, R.D., Sdrolias, M., Gaina, C., Roest, W.R., 2008, Age, spreading rates, and spreading asymmetry of the world’s ocean crust: Geochemistry, Geophysics, Geosystems, v. 9, Q04006, doi:10.1029/2007GC001743.

Oldow, J.S., 1984, Evolution of a late Mesozoic back-arc fold and thrust belt, northwestern Great Basin, U.S.A.: Tectonophysics, v. 102, p. 245-274.

Orr, P.C., 1956, Pleistocene man in Fishbone Cave, Pershing County, Nevada: Nevada State Museum Department of Archaeology Bulletin, v. 2, p. 211-220.

Passchier, C.W., and Truow, R.A.J., 2005, Microtectonics. Berlin: Springer-Verlag, 289 pp.

Paterson, S.R., and Vernon, R.H., 1995, Bursting the bubble of ballooning plutons; a return to nested diapirs emplaced by multiple processes: Geological Society of America Bulletin, v. 107, p. 1356-1380.

Profett, J.M., Jr., and Dilles, J.H., 1984, Geologic map of the Yerington District, Nevada: Nevada Bureau of Mines and Geology Map 77, scale 1:24 000, 1 sheet.

Quinn, M.J., Wright, J.E., and Wyld, S.J., 1997, Happy Creek igneous complex and tectonic evolution of the early Mesozoic arc in the Jackson Mountains, northwest Nevada: Geological Society of America Bulletin, v. 109 (4), p. 461-482.

Rai, V.N., 1969, Geology of a portion of the Nightingale and Truckee ranges, Washoe and Pershing counties, Nevada: [thesis].

Reheis, M.C., Sarna-Wojcicki, A.M., Reynolds, R.L., Repenning, C.A., and Mifflin, 59

M.D., 2002, Pliocene to Middle Pleistocene lakes in the Western Great Basin— Ages and connections, in Hershler, R., Madsen, D.B., and Currey, D.R., eds., Great Basin Aquatic Systems History: Smithsonian Contributions to the Earth Sciences 33, p. 53–108.

Russell, B.J., 1984, Mesozoic geology of the Jackson Mountains, northwestern Nevada: Geological Society of America Bulletin, v. 95, p. 313-323.

Russell, I.C., 1885, Geological History of Lake Lahontan, a Quaternary Lake of Northwestern Nevada: U.S. Geological Survey Monograph 11, 288 pp.

Saleeby, J.B., 2003, Segmentation of the Laramide Slab—evidence from the southern Sierra Nevada region: Geological Society of America Bulletin, v. 115, p. 655- 668.

Saltus, R.W., and Jachens, R.C., 1995, Gravity and basin-depth maps of the Basin and Range Province, Western United States: Geophysical Investigalions Map GP- 1012.

Sears, P.B., and Roosma, A., 1961, A Climatic sequence from two Nevada caves: American Journal of Science, v. 259, p. 669-678.

Silberling, N.J., and Wallace, R.E., 1969, Stratigraphy of the Star Peak Group (Triassic) and Overlying Lower Mesozoic Rocks, Humboldt Range, Nevada: U.S. Geological Survey Professional Paper 592, 50 pp.

Smith, J.G., McKee, E.H., Tatlock, D.B., Marvin, R.F., 1971, Mesozoic granitic rocks in northwestern Nevada: A link between the Sierra Nevada and Idaho Batholiths: GSA Bulletin, v. 82, p. 2933-2944.

Smith, W.C., and Guild, P.W., 1942, Tungsten deposits of the Nightingale district, Pershing County, Nevada: U.S. Geological Survey Bulletin 936-B.

Speed, R.C., 1978, Basinal terrane of the early Mesozoic marine province of the western Great Basin, in Howell, D.G., and McDougall, K.A., eds., Mesozoic Paleogeography of the Western United States: Pacific Section, Society of Economic Paleontologists and Mineralogists, Pacific Coast Paleogeography Symposium 2, p. 237-252.

Stager, H.K., and Tingley, J.V., 1988, Tungsten deposits of Nevada: Nevada Bureau of Mines and Geology Bulletin 105.

Stewart, J.H., 1998, Regional characteristics, tilt domains, and extensional history of the later Cenozoic Basin and Range Province, western North America: Special Paper – Geological Society of America, v. 323, p. 47-74.

Thole, R.H., and Prihar, D.W., 1998, Geologic map of the Eugene Mountains, northwest Nevada: Nevada Bureau of Mines and Geology Map 115.

Tikoff, B. and C. Teyssier, 1992, Crustal-scale, en echelon “P-shear” tensional bridges; a possible solution to the batholithic room problem, Geology, v. 20 (10), p. 927- 930.

United States Geological Survey, 1999, Map Accuracy Standards: Fact Sheet FS-171-99.

Van Buer, N.J., Miller, E.L., Dumitru, T.A., 2009, Early Tertiary paleogeologic map of the northern Sierra Nevada batholith and the northwestern Basin and Range: Geology, v. 37, p. 371-374.

Van Buer, N.J. and Miller, E.L., 2010, The Sahwave Batholith, NW Nevada: Cretaceous arc flare-up in a basinal terrane: Lithosphere, v. 2, p. 423-446.

Whitehill, C. S., 2009, Cenozoic evolution of the Shawave-Nightingale horst block, northwestern Basin and Range, Nevada, U.S.A. [thesis].

Willden, R., 1964, Geology and Mineral Deposits of Humboldt County, Nevada: Nevada Bureau of Mines and Geology Bulletin, v. 59, 154 pp.

Wooden, J.L., Kistler, R.W., Tosdal R.M., 1999, Strontium, lead, and oxygen isotopic data for granitoid and volcanic rocks from the northern Great Basin and Sierra Nevada, California, Nevada, and Utah, U.S. Geological Survey Open File Report 99-569, 20 pp.

Wyld, S.J., 1990, Paleozoic and Mesozoic rocks of the Pine Forest Range, northwest Nevada, and their relation to volcanic arc assemblages of the western U.S. Cordillera, in Harwood, D.S., and Miller, M.M., eds., Late Paleozoic and Early Mesozoic Paleogeographic Relations; Klamath Mountains, Sierra Nevada, and Related Terranes: Geological Society of America Special Paper 255, p. 219-237.

Wyld, S.J., 2000, Triassic evolution of the arc and backarc of northwest Nevada, and evidence for extensional tectonism, in Soreghan, M.J., and Gehrels, G.E., eds., 61

Paleozoic and Triassic Paleogeography and Tectonic Evolution of Western Nevada and Northern California: Boulder, Colorado, Geological Society of America Special Paper 347, p. 185-208.

Wyld, S.J., Rogers, J.W., and Wright, J.E., 2001, Structural evolution within the Luning- Fencemaker fold-thrust belt, Nevada: progression from back-arc basin closure to intra-arc shortening: Journal of Structural Geology, v. 23, p. 1971-1995.

Wyld, S.J., and Wright, J.E., 2001, New evidence for Cretaceous strike-slip faulting in the United States Cordillera and implications for terrane-displacement, deformation patterns, and plutonism: Amereican Journal of Science, v. 301, p. 150-181.

Wyld, S.J., and Wright, J.E., 2005, Early Cretaceous margin-parallel, dextral faulting and terrane translation in the U.S. Cordillera: Geological Society of America Abstracts with Programs, v. 37 (4), p. 102.

Zirkel, F., 1876, Microscopical Petrography: Professional Papers of the Engineer Department, U.S. Army, No. 18, 297 pp.

Zones, C.P., 1961, Ground-Water Reconnaissance of Winnemucca Lake Valley, Pershing and Washoe Counties, Nevada: U.S. Geological Society Water-Supply Paper 1539-C, 18 pp. 62

CHAPTER 3: THE SAHWAVE BATHOLITH, NW NEVADA: CRETACEOUS ARC FLARE-UP IN A BASINAL TERRANE

Nicholas J. Van Buer and Elizabeth L. Miller Department of Geological and Environmental Sciences, Stanford University 450 Serra Mall, Bldg. 320, Stanford, CA 94305-2115

A version of this chapter has been published in Lithosphere and copyright has been assigned to the Geological Society of America. GSA’s copyright policy explicitly permits reproduction in an author’s dissertation.

2010 Geological Society of America Van Buer, N.J. and Miller, E.L., 2010, The Sahwave Batholith, NW Nevada: Cretaceous arc flare-up in a basinal terrane: Lithosphere, v. 2, p. 423-446, doi: 10.1130/L105.1.

ABSTRACT

Detailed mapping and SHRIMP U-Pb geochronology centered around the

Nightingale and Sahwave Ranges, about 100 km northeast of Reno, Nevada, reveal that most of the Mesozoic basement in this area is composed of predominantly granodiorite- composition plutonic rocks intruded ca. 110–88.5 Ma. These rocks are similar in age, petrology, and composition to the mid-Cretaceous eastern part of the Sierra Nevada

Batholith, and are likely related. The youngest plutonic rocks, ca. 93–88.5 Ma, form a large, compositionally zoned intrusive suite, referred to as the Sahwave Intrusive Suite.

This suite is composed of a set of nested, inward-younging intrusions, varying between mafic, equigranular granodiorite around the periphery to more felsic, K-feldspar- megacrystic granodiorite in the center. The Sahwave Intrusive Suite is coeval with the

Cathedral Range intrusive event along the crest of the Sierra Nevada, including the

Tuolumne Intrusive Suite. The geochemistry and petrology of this intrusion also support similar magma genesis and emplacement.

Intrusions of the Cathedral Range intrusive event in the Sierra Nevada were emplaced along the margin of North American continental crust, whereas the Sahwave

Intrusive Suite was intruded into a thick package of basinal metasedimentary rocks that were likely underlain by transitional crust. More primitive initial 87Sr/86Sr and εNd values (ca. 0.7047 and –0.2, respectively) reflect this difference. In light of this likely fundamental difference in lower-crustal character, other factors, possibly related to subducted, water-rich material, must be responsible for creating similar melting conditions among the series of large intrusions that represent the last magmatic flare-up of the Cretaceous arc. 64

INTRODUCTION

The Mesozoic Sierra Nevada Batholith preserves an extensive record of continental-margin arc magmatism that serves as a classic, worldwide model, especially for high-intrusive-flux magmatism. Previously, however, only reconnaissance-level studies (e.g., Smith et al., 1971; Barton et al., 1988; Van Buer et al., 2009) have explored the possibility that this batholith might extend past the Sierra Nevada mountains into the

NW Basin and Range province (Fig. 3.1), where Mesozoic relationships are obscured by

Cenozoic volcanism and basin development related to extensional faulting.

Consequently, many published figures depicting the Sierra Nevada Batholith are truncated against the edge of the Basin and Range or the Nevada border (e.g. Tikoff and de Saint Blanquat, 1997; DeGraaff-Surpless et al., 2002; Lackey et al., 2005), and the

Sierra Nevada Batholith is often considered to be restricted to the mountains it was named for. However, boundaries as recent as the Neogene limit of Basin and Range extension (dotted line, Fig. 3.1), which defines the eastern scarp of the Sierra Nevada, would seem to rather arbitrarily delimit the much older Mesozoic Sierra Nevada

Batholith. Although Mesozoic outcrops in the Basin and Range are less continuous and more deeply weathered than those in the glacially-scoured Sierra Nevada, the Sahwave and Nightingale Ranges, about an hour NE of Reno, NV, form a broad, uplifted horst block of Mesozoic basement that is well suited for investigating the relationship between plutonic rocks in the NW Basin and Range and in the Sierra Nevada (Figs. 3.1,3.2).

Detailed mapping in the Sahwave and Nightingale Ranges, combined with reconnaissance of the surrounding areas, was used to identify distinct intrusive units for further quantitative study. Most of the intrusive units in this area were identified as 65

126° W 124° W 122° W 120° W 118°118° W 116° W 114° W 112° W 110° W 42° N OR Marginal Idaho Batholith ID 87 86 CA terranes Sr/ Sr = 0.706 S n a ain 44° N Klamath ke Rive r Pl 40° N Range NV

area of Study Fig. 2 42° N W a l k e r L a n e area

Sierra Nevada Batholith 38° N LFTB B a s i n UT a n d

R a n g e 40° N

36° N North

American 38° N AZ N craton 34° N

36° N

Peninsular 32° N Ranges 500 km Batholith 34° N

122° W 120° W 118° W 116° W 114° W 112° W 110° W 108° W Figure 3.1. Plutons of the Mesozoic magmatic arc (white) are most prominently exposed in the Idaho Batholith, the Sierra Nevada Batholith, and the Peninsular Ranges Batholith. The Sierra Nevada Batholith was emplaced across the boundary between the continental lithosphere of cratonal North America and a variety of marginal terranes which have oceanic- or transitional-affinity lithosphere as defined by the initial 87 Sr/ 86 Sr = 0.706 line (dashed line; from Farmer and DePaolo, 1983). The main part of the Basin and Range Province is outlined by a dot-dashed line. The Walker Lane accommodates right-lateral shear near the western boundary of the Basin and Range (Wesnousky et al., 2005). The Luning-Fencemaker Thrust Belt (LFTB) is developed in Mesozoic basinal sequences (Oldow, 1984). Distribution of Mesozoic intrusions modified from King and Beikman (1974).

Figure 3.2 (next page): Though disrupted by Basin and Range extension and largely buried by Ceneozoic cover (uncolored), Mesozoic intrusive rocks (medium grey) comprise most of the pre-Cenozoic rocks along a belt stretching NNE from the Lake Tahoe Region. Rocks of the Sahwave Intrusive Suite shown in very dark grey/black. 66

120º W 119º W 118º OREGON 42º N NEVADA

Explanation of Map Units 2004 Stanford seismic line

Cenozoic cover (Lerch et al., 2007) shaded by inferred basement Inner Sahwave intrusive suite

Outer Sahwave intrusive suite BASIN AND RANGE PROVINCE Unexposed Sahwave suite?

Other Mesozoic intrusions

Other pre-Cenozoic rocks 41º N Geologic map compiled mainly from Jennings et al., 1977; Irwin and Wooden, 2001; Bateman, 1992; John, 1983; Tikoff and de Saint Blanquat, 1997; Hirt, 2007; Saleeby et al., 2008. Ages are from many sources.

NEVADA

Selenite Range Susanville CALIFORNIA NVB 212 Bluewing Mts. area of Honey L. Fig. 3 NVB 286 Granite

Pyramid Lake Springs Valley

Nightingale R. Trinity Range Sahwave R.

Lake Range

40º N

B A S I N S I E R R A N E V A D A A N D Reno BASIN AND RANGE PROVINCE

Lake R A N G E Tahoe

Sacramento 67 belonging to a single, very large, roughly concentrically zoned intrusive suite, emplaced at about 90 Ma, referred to here as the Sahwave Intrusive Suite (Fig. 3.2). Zoned intrusive suites of approximately the same age in the Sierra Nevada, such as the

Tuolumne Intrusive Suite, have received detailed geochronological, mineralogical, geochemical, and structural study due to vigorous and ongoing debate about their petrogenesis and emplacement (e.g., Bateman, 1992; Coleman et al., 2004; Zak and

Paterson, 2005; Hirt et al., 2007; Gray et al., 2008), and therefore provide an excellent dataset for comparison with the Sahwave Intrusive Suite. As the first report of its kind in this region, this paper attempts to set forth several types of basic data, from map data and rock descriptions to modal mineralogy, U-Pb geochronology, and major, trace-element, and isotope geochemistry. Comparison of data between these intrusive suites allows us to evaluate whether the Sierra Nevada Batholith should be considered to extend into the

NW Basin and Range (Fig 3.1). Furthermore, differences between these regions of high intrusive flux may have important implications for arc flare-up models.

REGIONAL GEOLOGIC SETTING

Subduction-related arc magmatism in the Cordillera began in the Triassic and continued episodically into the Late Cretaceous (and into the Paleocene north of the

Snake River Plain and in southern Arizona; Fig. 3.1). The resulting batholithic belt has been variably disrupted by Cenozoic extension and translation and now forms several distinct segments, including the Idaho Batholith, the Sierra Nevada Batholith, and the

Peninsular Ranges Batholith (Fig. 3.1). The final episode of magmatism in California and Nevada spanned ca. 120-85 Ma, and was particularly voluminous during in the latter 68 half of this period (e.g., Barton et al., 1988; Ducea, 2001). In most of the U.S. Cordillera, the Cretaceous batholith exhibits a regular younging pattern from west to east that is generally mirrored by geochemical trends from more mafic to more felsic (e.g., Evernden and Kistler, 1970; Hyndman, 1983; Silver et al., 1979).

One of the most distinctive features of the Sierra Nevada Batholith is the series of large, compositionally zoned intrusions of the Cathedral Range intrusive event, such as the Tuolumne Intrusive Suite, emplaced along the eastern edge of the main Sierra Nevada

Batholith at the very end of Cretaceous arc magmatism between about 94–83 Ma

(Evernden and Kistler, 1970; Kistler et al., 1986; Tikoff and Teyssier, 1992).

Representing a high level of magmatic flux (e.g., Ducea, 2001), these intrusions generally exceed 1000 km2 in area, and are characterized by central megacrystic K-feldspar granites or granodiorites surrounded by more mafic equigranular granodiorites (e.g., Bateman,

1992; John and Robinson, 1982; Titus et al., 2005; Hirt, 2007; Saleeby et al., 2008).

Similar large, zoned intrusions are also present in the Peninsular Ranges Batholith (Fig.

3.1), although these are somewhat older (primarily ca. 99–92 Ma) and have tonalite and trondjhemite as well as granodiorite compositions (e. g., Gastil, 1983; Walawender et al.,

1990).

The Cretaceous Sierra Nevada and Peninsular Ranges Batholiths, which contain these large intrusions along their east sides, straddle the boundary between North

American continental crust and oceanic terranes to the west, as approximated by the initial 87Sr/86Sr = 0.706 line (Fig. 1; e.g., Gastil, 1975; Saleeby, 1981; Kistler, 1990); in contrast, the locus of Cretaceous magmatism between the Sierra Nevada and western

Idaho (Fig. 3.1) does not appear to be adjacent to regular continental crust. Wall rocks to 69 the Cretaceous intrusions in this area include a basinal terrane of lower Mesozoic deep marine strata and the early Mesozoic arc terranes bounding it to the northwest and southwest (Fig 3.2; e.g., Speed, 1978; Quinn et al., 1997; Wyld, 2000). These rocks have regionally been metamorphosed to subgreenschist to lower greenschist grade but often reach amphibolite grade proximal to Mesozoic intrusions (e.g., Willden, 1964; Bonham,

1969; Johnson, 1977; Barton et al., 1988). The basinal strata, which belong to the monotonous Late Triassic (Norian) to earliest Jurassic Auld Lang Syne Group, are essentially submarine fan deposits, metamorphosed into slate/phyllite with subordinate quartzite lenses and rare calc-silicate/marble layers (Burke and Silberling, 1973; Speed,

1978). Correlative, but thinner, strata overlie the shelfal, earlier Triassic Star Peak Group east of the main locus of Cretaceous magmatism (Silberling and Wallace, 1969), but further west the basinal strata exceed 6 km, with no base exposed (Compton, 1960; Burke and Silberling, 1973; Speed , 1978). Jurassic shortening associated with the Luning-

Fencemaker Thrust Belt (Fig. 3.1) has folded, thrust, and thickened this basinal sequence

(Oldow, 1984).

The metamorphic and plutonic rocks of the northern Sierra Nevada and the northwest Basin and Range are unconformably overlain by Eocene, Oligocene, and

Miocene volcanic and sedimentary rocks (Fig. 3.2). This widespread unconformity represents a profound change from erosion in the latest Cretaceous and early Tertiary to active deposition of volcanic and sedimentary strata in the Eocene to Miocene and is an important datum for reconstructing geologic relationships prior to Miocene extension and related tilting (Van Buer et al., 2009). Uplift and erosion of the Tertiary strata has resulted in exposure of the unconformity and underlying Mesozoic basement in the tilted 70 footwalls of most major Basin and Range normal faults, leaving a discontinuous

Mesozoic outcrop pattern (Figs. 3.1, 3.2).

Although the Cretaceous Cordilleran magmatic arc has been traced across NW

Nevada (Fig. 3.1) based on reported pluton ages between 105–85 Ma (e.g. Smith et al.,

1971; Barton et al., 1988; Wooden et al., 1999), reconnaissance studies and compilations have not adequately addressed the character of the intrusions across this intervening region. Previous mapping in northwestern Nevada includes thorough coverage only at

1:250,000 scale, which does not differentiate between separate plutonic units (Willden,

1964; Bonham, 1969; Willden and Speed, 1974; Johnson, 1977). More detailed work has been completed on Jurassic and Triassic intrusions in western Nevada (e.g. John et al.,

1994), and on intrusions in the gold-producing region of north-central and northeast

Nevada (Fig. 3.1; reviewed in du Bray, 2007). For Cretaceous plutons in NW Nevada, some structural and geochronological work has been completed (e.g., Wyld and Wright,

2001; Ciaverella and Wyld, 2008; Colgan et al., 2010), but no detailed petrologic or geochemical data have been published that are adequate to address the magma genesis of these intrusions and their relationship to contemporaneous intrusions in the Sierra

Nevada.

CRETACEOUS PLUTONIC ROCKS OF NORTHWEST NEVADA

Although intrusive rocks can be found scattered throughout much of the western

Basin and Range province, they comprise a majority of the pre-Cenozoic outcrop in an area trending NNE from the Lake Tahoe area across NW Nevada (Fig. 3.2; Barton et al.,

1988; Van Buer et al., 2009). Plutons in this area are not tightly stitched at the level of 71 exposure, but rather are often separated by substantial areas of metamorphic outcrop, often more than ten kilometers across (Fig. 3.2). Intrusive rocks include quartz monzonite and rare diorite/quartz diorite, but are predominantly granodiorite (cf. Smith et al., 1971). Published geochronology indicates intrusion during the Jurassic (ca. 200-

160 Ma) and the Cretaceous (ca. 115-85 Ma) with the greatest intrusion fraction between ca. 105-90 Ma, (Rai, 1969; Evernden and Kistler, 1970; Smith et al., 1971; Morton et al.,

1977; Marvin and Cole, 1978; Garside et al., 1992; John, 1992; Oldenburg, 1995; Wyld,

1996; Quinn et al., 1997; Wyld and Wright, 1997; Wooden et al., 1999; Wyld et al.,

2001, Van Buer and Wooden, 2007), although many intrusions remain undated or poorly- dated.

A particularly continuous area of Cretaceous plutonic outcrop occurs in the

Sahwave and Nightingale ranges, which together form a broadly synclinal horst, with major normal faults along the east side of the Sahwave Mountains and the west side of the Nightingale Range (Fig. 3.3). The bulk of both mountain ranges is granodiorite. This area of intrusive rock is separated from other plutons on the south and northeast by several kilometers of metamorphic wall rocks (Fig. 3.2), making the Sahwave and

Nightingale Ranges a well-bounded target for detailed study. However, because granodiorite outcrops in the Selenite Range, to the northwest, and the Trinity Range, to the east, are potentially contiguous, if not for intervening Cenozoic cover, these areas were also selected for reconnaissance study (Fig. 3.2). The nearest intrusive outcrops to the west, in the Lake Range, are visually dissimilar, and were not closely studied. In the

Sahwave and Nightingale Ranges (Fig. 3.3), a number of reports and theses include local, more detailed mapping (Smith and Guild, 1942; East and Trengrove, 1950; Rai, 1969; 72

elevation (km) 3

A’ JTrm Kjp Map Explanation

Tertiary

Ryholitic dikes and plugs

Volcanic and sedimentary strata Tu Andesitic to basaltic dikes

Diabase dikes

Cretaceous Sahwave intrusive suite Aplite, pegmatite, Sahwave Granodiorite Ks and leucogranite intrusions bend Granodiorite of Bob Spring Kbs (Units shown in lighter shade where

inferred to underlie valley sediments) Kbs School Bus Granodiorite Ksb

Granodiorite of Juniper Pass Kjp Diorite intrusions

Earlier Mesozoic 10 km

Selenite Granodiorite Kse

Power Line intrusive complex Kpl

Auld Lang Syne Group JTrm basinal metasedimentary rocks

Symbols no vertical exaggeration

Contact Faults: Inferred Concealed (Balls on downthrown side) Gradational Ks contact Attitudes: Magmatic Metamorphic Bedding Concealed foliation foliation contact Sample locality Mineral lineation Kbs

Figure 3.3 (continued on next page): Detailed map of the Sahwave and Nightingale Ranges (location shown with box in Fig. 3.2), showing distribution of Cretaceous plutonic bodies,

older wall rocks, and Cenozoic cover. Based on mapping at Kjp 1:24,000 scale but shown here at 1:180,000 scale. JTrm Tu

A 3 elevation (km) 0 73

Kse Figure 3.3 (continued) Kjp JTrm A’ K U M I V A BLUEWING MTS. N JTrm Kjp Tu NVB-206

V A L L E Y Tu ?

10 km

Kjp Kjp 1:180,000 scale ? Tu 40° 15’ N Tu

Kpl Kbs

Kjp Kbs Kbs Kpl Tu

Kpl Kjp

S A G E H E N V A L L E Y Kpl NVB-208

Kbs Tu JC03-SV3 NVB-207 Kpl

JTrm Kjp Kbs Ks

Ksb G R A N I T E S P R I N G S V A L L E Y E L L A V S G N I R P S E T I N A R G

JTrm Tu Tu Kjp Kjp S A H W A V E R A N G E Tu

JTrm JTrm

Tu 40° N

N I G H T I N G A L E JTrm R A N G E Tu Tu 119° W 119° 15’ W Tu A Tu 74

Fanning, 1982; Stager and Tingley, 1988; Whitehill, 2009). Most of these papers relate to exploration of the tungsten-mining district along the southwest margin of the

Cretaceous intrusive contact, in the southern Nightingale Range (Fig. 3.3), and contain very little data pertaining to the igneous rocks themselves.

NEW MAPPING IN THE SAHWAVE AND NIGHTINGALE RANGES

Mapping of the Sahwave and Nightingale Ranges was completed at 1:24,000 scale (Plates 1–8; reduced to 1:180,000 in Fig. 3.3). Each mountain range is generally more rugged on the side between its crest and its bounding normal fault, and the Sahwave

Range, in particular, becomes higher and rockier to the north. However, even the best outcrops in this area are patchy and deeply weathered as compared to the continuous outcrops of the Sierra Nevada crest. The southern Sahwave Range (Fig. 3.3) forms a particularly low-relief upland, characterized by a thick blanket of grus, which nourishes the range’s namesake sagebrush (northern Paiute sai’-wav; Fowler and Fowler, 1971).

Map units fall into three basic categories: the overlying Cenozoic strata, early Mesozoic metamorphic wall rocks, and Cretaceous intrusive rocks. The intrusive rocks are further subdivied into two groups: units that frequently have gradational contacts with each other, are more or less concentrically arranged, and have magmatic foliation that is generally weak, absent, or contact-parallel, and are apparently cogenetic, are herein informally named the Sahwave Intrusive Suite; whereas units with strong, roughly north- south foliation, sharply crosscut by members of the aforementioned suite, are considered to be distinct pre-existing intrusions.

Cenozoic Strata

Oligocene and Miocene volcanic and sedimentary rocks unconformably overlie all Mesozoic units (Fig. 3.3). Together with recent alluvial deposits, these rocks fill part of the area between the Sahwave and Nightingale Ranges, as well as Cenozoic extensional basins to the east, west, and north of the Sahwave-Nightingale horst (Figs.

3.2, 3.3). The volcanic units range from basalt flows to silicic ignimbrites, and the interbedded sediments (not thoroughly lithified yet) range from landslide deposits and fanglomerates to lacustrine clays (cf. Whitehill, 2009). Additionally, the study area is cut by a few generations of dikes, ranging from rhyolitic to basaltic in composition, which also tend to be more resistant to erosion than the surrounding rock (Fig. 3.3). The coarsest are diabase (with chilled margins), but most have an aphanitic matrix. Although these dikes have not been dated, their fine-grained nature suggests that they are substantially post-magmatic, and they are compositionally similar to volcanic packages in the overlying Tertiary strata.

Country Rocks

The wall rocks of the Sahwave batholith are mostly metamorphosed mudstone/shale with interbedded sandstone layers and lenses. A few discontinuous, 10–

100 m thick, coarsely crystalline marble layers are present south of the batholith, but calcareous layers are rare in the metamorphic rocks to the north (Fig. 3.3). These rocks have been identified as belonging to the Triassic to lower Jurassic Auld Lang Syne Group

(Johnson, 1977). Away from the batholith, metamorphic grade is subgreenschist to lower greenschist, and original bedding is clearly seen. Fold axes and foliation in the adjacent 76

Bluewing Mountains (Fig. 3.3; Plate 2) trend NE-SW, and exhibit top-to-the-SE vergence, consistent with Jurassic deformation in nearby parts of the Luning-Fencemaker thrust belt (Fig. 3.1; Oldow, 1984). Adjacent to the batholith, Triassic/Jurassic strata are metamorphosed to siliceous hornfels or biotite schists, and bedding is often tightly to isoclinally folded with a sub-vertical axial-planar foliation that is broadly parallel to the intrusive contact (Fig. 3.4A). A strong sub-vertical mineral lineation is also present within ~100m of the intrusive contact, although it is often obscured by a subparallel intersection lineation. In the Bluewing Mountains, along the northern edge of the batholith, the zone of contact-parallel foliation is only a few hundred meters wide, whereas to the southwest, in the Nightingale mining district, foliation is sub-parallel to the contact over the entire exposed outcrop, up to 5 km away from the intrusion (Fig. 3.3;

Plates 2, 6). This NW-SE foliation is anomalous compared to NE-SW Jurassic structural trends which tend to dominate in surrounding areas (e.g., Oldow, 1984).

With some exceptions, the generally quartzofeldspathic composition of the metasedimentary rocks is not conducive to the growth of diagnostic minerals besides white mica and occasional biotite. The Nightingale mining district contains a number of skarn deposits in the contact aureole of the batholith where calcareous layers have been metamorphosed, yielding grossular/almandine, clinozoisite/epidote, and more rarely tremolite, diopside, and scheelite, in addition to the standard quartz, ±albite, and calcite.

White mica pseudomorphs, apparently after both andalusite (square rods) and also cordierite (dark, mouse-dropping shapes), can be found in some of the more pelitic layers in the Bluewing Mountains near the northern margin of the batholith (Fig. 3.3).

However, large (to over 5 cm) andalusite crystals remain intact in at least one area about 77

Figure 3.4 (next two pages): (A). Looking down at tight folds in interbedded shales (dark) and calcareous siltstone layers (light), ~ 100 m from intrusive contact in southern Nightingale Range. Deformation presumed Cretaceous since fold axes and lineations are aligned downwards paralleling the contact. Hammer for scale. (B). Thin section of Power Line Intrusive Complex under crossed polars. Note recrystallized biotite in center, strung out along a wavy foliation plane between feldspar and recrystallized quartz grains. For all thin section images, B = biotite, H = hornblende, K = potassium feldspar, M = magnetite, P = plagioclase, Q = quartz, S = sphene. (C). Hand sample of the Grano- diorite of Juniper Pass. Dark grains are biotite and hornblende. Honey-colored grains are sphene, e.g., near the top left corner. (D). Thin section of Granodiorite of Juniper Pass under crossed polars. Note equigranular texture with biotite (upper right), and hornblende (lower left), in addition to microcline, plagioclase, and quartz. (E). Complex magmatic structures in the Granodiorite of Juniper Pass, ~2 km from outer margin. Note the wavy compositional layering, especially just below left of the hammer. The central, more leucocratic dike also displays complex interfingering with a more mafic phase just left of the hammer handle. Biotite within the more lighter phases (black dots) is coarser than in the darker phases. (F). Outcrop of Granodiorite of Juniper Pass ~ 1 km from outer margin, showing elongated mafic enclaves and magmatic foliation (parallel to black line). Hammer head for scale at top of rock. (G). Thin section of Granodiorite of Bob Spring under crossed polars. Note the large, poikilitic K-feldspar, and the chloritized biotite (dark) at lower left. (H). Hand sample of Sahwave Granodiorite, showing large K-feldspar phenocryst within a more equigranular matrix. (caption contiued on second page) 78

A B

1 mm

C 1 cm D

1 mm E F

G H

1 mm 1 cm Figure 3.4 (continued) 79

Q P B M B B S Q

B P K

P S

B P Q P

B K P S

Figure 3.4 (continued): (I). Thin section of Sahwave Granodiorite under crossed polars, showing conspicuous sphene wedges and small, ragged biotites. (J). Megacryst- rich pods in the Sahwave Granodiorite, outlined and labeled “m”, surrounded by relatively leucocratic material. Arrow points to hammer for scale. (K). Thin section of School Bus Granodiorite under crossed polars, showing large biotite (left), conspicuous sphene (center right), and myrmekitic contact between quartz and feldspar (upper left). (L). Vertical contact between School Bus Granodiorite (right half) and Power Line Complex (left half). Finger for scale. (M). Mingling and mixing relations between diorite and Granodiorite of Juniper Pass. Note interfingering of darker and lighter units, as well as continuous gradations in color index. The dark splotches at lower right are lichen. Mechanical pencil for scale. (N). Aplite dike. Composite layering can be seen dipping steeply to the left (west). Hammer for scale at very top of rock. 80

1 km from the northern contact, growing in random orientations that cut across the foliation.

Early Intrusive Units

The oldest intrusive unit, informally referred to as the Power Line intrusive complex (Kpl), occupies the northwestern Nightingale Range (Fig. 3.3), and is predominantly a medium-grained biotite hornblende granodiorite with 5–10 mm K- feldspar phenocrysts. However, this unit also includes many unmapped dikes and pods of darker granodiorite and diorite ranging from centimeters to hundreds of meters in dimension. Some of these are fine grained, weathering to a blue-grayish color, but all subunits share a similar, generally north-south oriented, steeply-dipping solid-state foliation (Fig. 3.3; Plates 3, 4). This strong foliation distinguishes the Power Line

Complex from all other intrusive units, including the Sahwave Intrusive Suite, which intrudes the complex and crosscuts its foliation. Although many of the finer-grained mafic enclaves appear to demonstrate magma mingling, relationships among these subunits are somewhat obscured by poor outcrop and the solid-state foliation. In thin section, the foliation is defined by biotite strung out along wavy foliation planes, and the sense of shear, if any, is unclear, as the rock bears no discernible lineation (Fig. 3.4B).

Biotite and quartz appear to have been largely recrystallized (Fig. 3.4B), but feldspars remain intact, displaying distinct undulatory extinction, suggesting solid-state deformation at temperatures of ca. 400–450 °C or warmer, depending on strain rate.

This unit also contains many large inclusions of metamorphic rock, mostly 5–200 m in length but including a 4-km-long potential roof pendant as well (Fig. 3.3; Plates 3, 4). 81

These are generally elongated in map view, and aligned subparallel to the foliation of the

Power Line Complex (Fig. 3; Plates 3, 4).

In the very northwestern corner of the study area and throughout the southern

Selenite range (Figs. 3.2, 3.3; Plate 1) is a distinct granodiorite here referred to as the

Selenite Granodiorite (Kse) after the “Selenite pluton” of Smith et al. (1971). This unit has a conspicuous, generally north-south magmatic foliation defined by the alignment of euhedral plagioclase and hornblende phenocrysts in rock with a hypidiomorphic igneous texture. Polysynthetic twinning in the plagioclase is frequently visible to the unaided eye. This unit is tentatively not included in the Sahwave Intrusive Suite, which intrudes it along a sharp contact (Fig. 3.3; Plate 1) and only rarely contains euhedral plagioclase.

THE SAHWAVE INTRUSIVE SUITE

The metamorphic rocks, the Power Line Complex, and the Selenite Granodiorite are intruded by members of the Sahwave Intrusive Suite, which consists of three concentric, partially intergradational intrusive units centered on the Sahwave Range and a distinct lobe-forming unit that stretches across the central Nightingale Range (Fig. 3.3).

Rocks of similar appearance also occur in the western Trinity Range, separated from the

Sahwave Range by Cenozoic fill in Granite Springs Valley, suggesting the Sahwave

Intrusive Suite may underlie much of this broad area as well (Fig. 3.2). The outermost and oldest intrusive unit is a medium-to-coarse-grained equigranular biotite hornblende granodiorite referred to as the Granodiorite of Juniper Pass (Kjp; Fig. 3.3). This unit is discernible by its conspicuous 4–8 mm biotite crystals. Additionally, large hornblende phenocrysts are common around the periphery of this intrusion, giving the rock a 82 characteristic “dalmatian” appearance (Fig. 3.4C, D). Hornblende and sphene are both present throughout the Sahwave Intrusive Suite, but only in the Granodiorite of Juniper

Pass does the hornblende form crystals notably larger than the 1–3 mm euhedral sphene.

In detail, the mineral proportions and color index of this unit varies quite a bit; in places it can be classified as a tonalite or a quartz diorite. Gradational compositional variation can sometimes be seen across large outcrops; more rarely, internal contacts can be discerned where slightly lighter and darker phases occur together. In a few places, straight or wavy compositional layers 1 cm – 1 m thick are bounded by sharp contacts (Fig. 3.4E). Many of these internal structures are subtle, and only readily seen in fresh outcrop, so it is possible that they are fairly pervasive. Mafic enclaves are found throughout the unit, but are only common within 1–2 km of the exterior contact. Enclaves are typically 5-30 cm in length and flattened by a ratio of 2:1 to 5:1 or more (Fig. 3.4F). Mafic schlieren are common in the same region. The Granodiorite of Juniper Pass has a discernible magmatic foliation that is defined by the alignment of mafic minerals and sometimes subhedral plagioclase, which is generally similar to the alignments of mafic schlieren and mafic enclaves as well (Fig. 3.4F). Magmatic foliation tends to be strongest near the outer contact, which it often parallels (Fig. 3.3).

The Granodiorite of Juniper Pass grades inward to the more felsic and uniform

Granodiorite of Bob Spring (Kbs), a medium-grained biotite granodiorite or granite, characterized by seriate K-feldspar phenocrysts up to ~ 2 cm. Although relative age relations with the Granodiorite of Juniper Pass are difficult to determine from the gradational intrusive contact, in map pattern the Granodiorite of Bob Spring appears to cut out the center of the Juniper Pass (Fig. 3.3) and is presumed to be younger. In the 83 field, this gradational contact is arbitrarily mapped where large K-feldspar phenocrysts become more conspicuous than large biotite crystals. Biotite in the Granodiorite of Bob

Spring is more homogeneously distributed, and generally no larger than 1 mm. The K- feldspar phenocrysts are poikilitic, mostly surrounding plagioclase and biotite (Fig.

3.4G), and are occasionally sieve textured and difficult to see. In general, Kbs is finer grained towards its center, and K-feldspar phenocrysts are less common. The

Granodiorite of Bob Spring bears equant quartz grains that are generally only about 1 mm in size but reach 3–5 mm in the southern part. Mafic minerals are often badly chloritized and feldspars show signs of sericitization. Foliation in this unit is usually absent or at least too weakly defined to measure.

The Sahwave Granodiorite (Ks), a K-feldspar-megacrystic biotite granodiorite

(Fig. 3.4H,I), intrudes the central part of the Granodiorite of Bob Spring along a generally shallowly dipping contact that is sharp on the north side but gradational along its south side (Fig. 3.3; Plates 5, 8). K-feldspar megacrysts are 2-4 cm across, somewhat poikilitic, and more abundant (usually 1–5% by volume) than in the Granodiorite of Bob Spring.

The abundance of K-feldspar megacrysts can vary greatly from place to place and at outcrop scale it is not uncommon to see distinct stringers and pods enriched in K-feldspar megacrysts, rarely up to as much as ~20% (Fig. 3.4J). The Sahwave Granodiorite forms relatively bold outcrops compared to adjacent parts of the Granodiorite of Bob Spring, but the rock is uniformly crumbly and often spheroidally weathered.

The Nightingale Range contains a distinct lobate unit referred to as the School

Bus Granodiorite (Ksb; Fig. 3.3). This unit is a relatively leucocratic granodiorite, distinguished by scattered 1–2 cm K-feldspar phenocrysts and 3–6 mm biotite flakes (Fig 84

3.4K). Unlike the main part of the Sahwave batholith, this lobe does not appear to be any more mafic around its outer edge, and is, in fact, remarkably homogeneous. Magmatic foliation is not generally distinguishable. The School Bus Granodiorite intrudes both the

Power Line Complex and the Granodiorite of Juniper Pass along sharp, vertical contacts

(Fig. 3.4L) that are fairly irregular at the map scale (Fig. 3.3; Plates 3, 4, 6, 7). Where it intrudes the Power Line Complex, the units are often separated by metamorphic screens and blobs 20-200 m thick (Fig. 3.3; Plates 3, 4).

Minor Intrusives

The southern part of the Granodiorite of Juniper Pass, contains a number of diorite/quartz diorite bodies, varying from tens of meters to over a kilometer in scale (Fig.

3.3; Plate 7). These fine- to medium- grained intrusions frequently contain ~5 mm euhedral plagioclase phenocrysts, and sometimes acicular hornblende crystals as well.

These diorite bodies appear to be coeval with the Sahwave batholith, often showing magma mingling and mixing structures such as lobate and interfingering contacts, streaky fine-scale intermingling, and outcrop-scale continuous compositional variation indicative of wholesale mixing (Fig. 3.4M).

Additionally, the Sahwave batholith and its country rocks are pervaded by a series of leucocratic dikes and sills that tend to be more resistant to weathering than the surrounding country rocks (Fig. 3.3). Most of these dikes demonstrate wide variations in grain size between aplite and pegmatite textures, often showing evidence for repeated intrusion (Fig. 3.4N). The pegmatites are generally muscovite bearing, and may also contain tourmaline (schorl) and rarely garnet. The dikes range from 1 cm to 100 m in 85 thickness and generally strike north-south (Fig. 3.3). Dips are often moderately shallow, but only locally consistent in direction of dip. A notable concentration of leucocratic dikes exists in the Nightingale Range, intruding the Power Line Complex (Fig. 3.3; Plate

3). These dikes crosscut the solid-state foliation of the Power Line Complex and are occasionally composite, containing a phase with scattered large K-feldspar and biotite phenocrysts, suggesting that they may be genetically related to the School Bus

Granodiorite, which could underlie this area at an unexposed level. Pegmatite dikes cutting the metamorphic rocks along the margins of the batholith are frequently folded and boudinaged in the foliation that is subparallel to the country-rock contact. A few broader leucogranite intrusions, which are more uniform in grain size and contain minor biotite, are present near the southern margin of the batholith (Fig. 3.3; Plates 6, 7, 8).

Intrusive Contacts

The intrusive contacts are generally not exposed well enough, or, when gradational, defined well enough to measure their attitudes directly, and furthermore are generally too irregular where exposed on the outcrop scale to make meaningful map-scale measurements directly. Contact attitudes, such as those shown on the cross section in

Figure 3.3, have been estimated from map patterns using three-point constraints in areas where the contact appears to be approximately planar. Where contact orientation is evident, it tends to be steeply dipping and subparallel to magmatic foliation, but there are a couple of notable exceptions. These are the contacts along the two largest metamorphic blocks or pendants at the southern end of the Sahwave Range, and the shallow contact where the Granodiorite of Juniper Pass underlies the Power Line Complex in the 86 northwestern Nightingale Range (Fig. 3.3; Plate 3, 4, 7, 8). It is not clear, however, whether these cases represent the true roof of the intrusion. In the northwestern

Nightingale range, the low-angle portion of the contact terminates westward as the top contact of a horizontal dike of the Granodiorite of Juniper Pass intruded into the Power

Line Complex (Fig. 3.3; Plate 3), suggesting that the contact in this area may simply surround a flap of wall rock that was in the process of being stoped off. External contacts of the Sahwave Intrusive Suite frequently dike into the metamorphic rocks and apparently surround stoped blocks (Fig. 3.3), suggesting that stoping is at least a locally important process. In other areas, external contacts are sometimes quite planar, demonstrating smooth curves that parallel foliation in the adjacent, subvertically lineated wall rocks (Fig. 3.3), suggesting that the wall rocks were flattened in pure shear and flowed ductilely downwards to accommodate the laterally expanding pluton.

CHRONOLOGY OF EMPLACEMENT

Although relative ages for the plutons can be determined from contact relations, the only published K/Ar hornblende (and biotite) ages are 91 ± 6 (88 ± 4) Ma and 95 ± 6

(92 ± 4 ) Ma for the Granodiorite of Juniper Pass and the Selenite Granodiorite, respectively (Smith et al., 1971). These error bars are about as large as the total span of ages. To more precisely define the timing and duration of magmatism in the study area, samples from each of the six main intrusive units were selected for age determination.

An additional sample of granodiorite from the Trinity Range, resembling the School Bus

Granodiorite, was dated to investigate whether the Sahwave Intrusive Suite might continue this far east (Fig. 3.2). 87

U-Pb SHRIMP Methods

Zircons from these seven samples (Table 3.1) were analyzed by secondary-ion mass spectrometry using the Stanford-U.S.G.S. SHRIMP-RG to yield U-Pb age determinations. Zircons were separated from each sample using standard procedures.

Sample zircons and chips of R33 standard zircons were mounted in epoxy, ground halfway through the grains with fine sandpaper, and polished with diamond compound.

All grains were imaged both in reflected light with an optical microscope and in cathodoluminescence (CL) using a JEOL 5600 scanning electron microscope to reveal zonation as well as cracks, inclusions, and other potential problem areas. U, Th, and Pb isotopes, along with Zr, Hf, La, Ce, Nd, Sm, Eu, Gd, Dy, Er, and Yb were analyzed with the Stanford/U.S.G.S. SHRIMP-RG (reverse geometry) using an oxygen ion beam between 4–6 nA and a spot size of 20–30 µm. Isotope ratios were normalized using zircon age standard R33 (419 Ma, Black et al., 2004) and concentration standard CZ3.

Age data was reduced using SQUID and ISOPLOT software (Ludwig, 2001, 2003) to yield 207Pb-corrected 206Pb/238U weighted-average ages (Table 3.1; Fig. 3.5). Complete data tables can be found in Appendix A1.

U-Pb SHRIMP Results

Zircons generally show crisp magmatic oscillatory zonation under CL and do not contain distinct cores (Figure 3.6). Only one grain, from NVB-286, appeared to have a distinct core and rim, but both parts gave exactly the same age. Individual grain analyses showed a moderately large amount of scatter, although most analyses spread out along or just above concordia (Fig. 3.5). Select analyses were dismissed (open symbols) because 88

TABLE 3.1. SHRIMP U-Pb GEOCHRONOLOGY SAMPLE DATA

Unit Sample Latitude Longitude Age* number (°N) (°W) (Ma)

Sahwave Granodiorite JC03-SV3 40° 07’ 56” 119° 04’ 05” 88.5 ± 2.0 Schoolbus Granodiorite NVB-207 40° 07’ 37” 119° 16’ 06” 91.2 ± 1.2 Granodiorite of Juniper Pass NVB-206 40° 18’ 47” 119° 01’ 04” 92.7 ± 1.4 Granodiorite in Trinity Range NVB-286 40° 10’ 13” 118° 46’ 20” 90.3 ± 0.6 Selenite Granodiorite NVB-212 40° 25’ 58” 119° 16’ 09” 96.3 ± 0.8 Power Line Intrusive Complex NVB-208 40° 08’ 38” 119° 13’ 20” 104.9 ± 0.8 boudinaged pegmatite dike NVB-280 40° 02’ 49” 119° 17’ 18” 89-97†

*Reported ages are 207Pb corrected 206Pb/238U weighted average ages with 2σ error. †This sample had no coherent age group. 89

238 206 207 206 238 U/ Pb Pb corrected Pb/ U ages 60 65 70 75 80 1 0 0 .07 9 5 .06

9 0 .05 .04 8 5 .03 JC03-SV3 Ks

100 .05

.04 95 .03 90 .02 NVB-207 Ksb

.06 9 5 .05 9 0 .04 Pb to 44 Ma 206 NVB-206 Kjp .03

to 139 Ma 207Pb/206PbPb/ Age (Ma) Age 9 5 207Pb/206Pb 207 .07

9 0 .06

.05 8 5 NVB-286 Trinity Range .06

100 .05

95 .04 NVB-212 Kse .03

110 .05

105 .04

.03 100 NVB-208 Kpl .02

9 5 .05

9 0 .04 SH-21 Kbs .03 8 5 2σ errors shown in both types of plot. 110 100 90 80 Age on concordia (Ma) Figure 3.5. SHRIMP U-Pb results. At left, selected spot ages used for weighted-mean ages are shown by solid bars. Rejected spot ages are shown by empty bars. Weighted averages are shown by gray lines. Diagrams on right are inverse concordia plots, showing accepted spot analyses with solid symbols and rejected spot analyses with empty symbols. Ages along concordia (gray line with tick marks) are shown on the corresponding tick marks at bottom. All errors are 2σ. 90

105.3 ±1.4

108.7±1.3

99.5 ±1.4 105.0 ±1.2

Figure 3.6. Cathodoluminescence images of representative zircons from the Power Line Intrusive Complex (NVB-208), which contains the most antecrystic zircons of any unit. All zircons demonstrate magmatic oscillatory zoning; younger and older zircons do not show systematic differences. 91 of discordance, high common lead, and lead loss in high-U zircons (Fig. 3.5; Table A1).

It is difficult to tell whether the spread in ages is caused by disturbed U-Pb systematics or actually represents prolonged periods of crystallization in a large active magma chamber episodically fed by new batches of magma. Older and younger zircons from individual samples are not visually different or distinguishable in CL images (example shown in

Fig. 3.6). Some of the significantly older ages can be ascribed to scavenging from slightly older plutonic rocks, such as the distinct population at 109.7 ± 0.8 Ma in NVB-

208 (Fig. 3.5). However, these rocks lack clear evidence of older inherited zircons (only one grain out of 77 analyzed was more than ~ 8 Ma older than the enclosing host rock, at a modest 139 Ma). The lack of significant inheritance suggests that these magmas may have originated at zircon-undersaturated conditions.

SHRIMP U-Pb results (Table 3.1; Figure 3.5) give ages for individual units in agreement with the relative ages inferred from intrusive relations. The Granodiorite of

Juniper Pass and the Granodiorite of Bob Spring give indistinguishable ages, but the latter is presumed to be younger from crosscutting map relations. These ages are also equivalent within error to published K/Ar hornblende and biotite ages (Smith et al.,

1971). Ages from the Sahwave Intrusive Suite span from ca. 93–88.5 Ma, demonstrating that this batholith is contemporaneous with the large intrusions of the ca. 95–83 Ma

Cathedral Range intrusive epoch defined in the Sierra Nevada Batholith (Evernden and

Kistler, 1970; Kistler, 1999). The sample of granodiorite from the western Trinity Range

(NVB-286) is also shown to have crystallized in this time range, at 90.3 ± 0.6 Ma, supporting the idea that it is part of the Sahwave Intrusive Suite and that these rocks may underlie much of the intervening Granite Springs Valley as well (Fig. 3.2). Whereas ages 92 associated with the Sahwave Intrusive Suite are clustered relatively tightly, spanning about four million years, the Selenite Granodiorite and the Powerline Complex are significantly older, at 96.3 ± 0.8 and 104.9 ± 0.8 Ma, respectively, justifying their classification as distinct units.

CONTINUITY OF THE CRETACEOUS CORDILLERAN BATHOLITH

Our initial study of batholithic rocks in the area around the Sahwave and

Nightingale Ranges in NW Basin and Range strongly supports the suggestion of Smith et al. (1971) and Barton et al. (1988) that the Cretaceous Cordilleran batholith is continuous across NW Nevada (Figs. 3.1, 3.2). Although obscured by Cenozoic cover, especially under the unbroken volcanic plateau covering NE California, SE Oregon, and part of NW

Nevada, Cretaceous batholithic rocks form a majority of Mesozoic outcrops along a

NNE-trending belt of the northwestern Basin and Range (Fig. 3.2; Barton et al., 1988;

Van Buer et al., 2009). The extent of this intrusive belt to the north and west is unclear due to complete Cenozoic volcanic cover, but relatively low upper-crustal seismic velocities compatible with granitoid rocks persist almost to the NW corner of Nevada

(Fig. 3.2; Lerch et al., 2007). Granodiorite units in the Sahwave and Nightingale area are similar to many described units in the Sierra Nevada Batholith, including units which, for example, contain conspicuous sphene, euhedral biotite and hornblende, or K-feldspar megacrysts. Large, concentrically-zoned intrusions are also common in the Sierra

Nevada (e.g., Bateman, 1992). U-Pb SHRIMP dating in the Sahwave and Nightingale area confirms earlier geochronologic estimation of Late Cretaceous ages simultaneous with major intrusion in the Sierra Nevada. Ages spanning from ca. 110 Ma (represented 93 by inherited zircons in the Power Line Complex) to ca. 88.5 Ma indicate a long-lived history of repeated intrusion in this part of the batholith, consistent with prolonged histories of magmatism in similarly sized areas of the Sierra Nevada Batholith

(e.g.,Bateman, 1992; Irwin and Wooden, 2001; Saleeby, 2008). These lines of evidence all support the idea that Cretaceous intrusive rocks in the study area formed in a broadly similar arc environment as those in the Sierra Nevada, and represent a continuation of the

Cretaceous Cordilleran arc across the NW Basin and Range (Fig. 3.1).

Whether or not the Cretaceous intrusions in NW Nevada should actually be considered to be part of the Sierra Nevada Batholith is largely a semantic issue. But if the boundaries of this Mesozoic batholith are to be set based on Mesozoic features, we note that the mostly Late Cretaceous intrusions of our study area lie due east of Lower

Cretaceous intrusions near Susanville, in the northernmost Sierra Nevada, which are generally considered to be part of the Sierra Nevada Batholith (Fig. 3.2; Oldenburg,

1995). Before Tertiary extension and translation across the Walker Lane (which is considered to be < 30 km at this latitude; Faulds, 2005) these two areas would have been even closer (Van Buer et al., 2009), representing the east and west edges of the eastward- younging batholith (Figs. 3.1, 3.2). Therefore, we tentatively suggest that the Cretaceous intrusions of NW Nevada be referred to as part of the Sierra Nevada Batholith. To better clarify the relationship between intrusions of NW Nevada and the Sierra Nevada, however, we analyze the magma genesis of the Sahwave Intrusive suite using detailed mineralogical and geochemical data, which we compare to similar data from the most well-studied intrusion of the same age in the Sierra Nevada, the Tuolumne Intrusive

Suite. 94

MINERALOGY AND GEOCHEMISTRY OF THE SAHWAVE INTRUSIVE

SUITE

Because the main, concentric part of the Sahwave Intrusive Suite appears to be younging-inwards, with mafic units grading into more felsic units, the rocks along a radial transect effectively record the magmatic evolution of the system over its four million year intrusive history. For this reason, the Sahwave Intrusive Suite was sampled from center to margin along a transect extending north from the central Sahwave

Granodiorite to the outer edge of the batholith and along a second, smaller transect through the School Bus lobe in the Nightingale Range (rows of black dots, Fig. 3.3).

Each transect contains samples spaced approximately a kilometer apart (Table 3.2), chosen from the most pristine outcrops available. Along these transects, the mineralogy records changes in the crystallizing assemblage and determines rock classification under the IUGS scheme (Streckeisen, 1976). Major and trace element chemistry respond in detail to element partitioning and mixing during melting, crystal-liquid fractionation, assimilation and other petrogenetic processes (e.g., Hildreth and Moorbath, 1988). Sr and Nd isotope systems are affected by radiogenic decay of Rb and Sm isotopes, and are therefore sensitive to the timing and extent of differentiation in the source region.

Together, these data provide information on the magma genesis of the system, and are appropriate for detailed comparison to similarly-sampled coeval intrusions in the Sierra

Nevada, such as the Tuolumne Intrusive Suite (e.g., Bateman, 1992; Hirt, 2007; Gray et al, 2008).

TABLE 3.2. MINERALOGY AND MAJOR ELEMENT CHEMISTRY

Sample data Modal mineralogy Sample Latitude Longitude Points Quartz K-feldspar Plagioclase Mafic Unit number (°N) (°W) counted (vol. %) (vol. %) (vol. %) (vol. %)

Main Sahwave transect Kjp SH-1 40° 19’ 49” 119° 00’ 37” 1248 23.6 2.6 48.6 25.2 Kjp SH-2 40° 19’ 21” 119° 00’ 46” 1058 19.8 6.4 54.8 18.9 Kjp NVB-206 40° 18’ 47” 119° 01’ 04” 1093 18.7 9.1 52.7 19.5 Kjp SH-5 40° 18’ 29” 119° 01’ 38” 1201 21.6 14.2 50.0 14.2 Kjp SH-6 40° 18’ 01” 119° 01’ 58” 1051 22.4 11.7 51.8 14.2 Kjp SH-7 40° 17’ 32” 119° 02’ 28” 1061 16.8 11.4 53.3 18.5 Kjp SH-8 40° 17’ 06” 119° 02’ 53” 856 17.4 2.8 59.7 20.1 Kjp SH-9 40° 16’ 39” 119° 03’ 22” 1051 15.8 4.2 60.4 19.6 Kjp SH-10 40° 16’ 10” 119° 03’ 32” 208 18.3† 13.0† 48.6† 20.2† Kjp SH-11 40° 15’ 38” 119° 03’ 48” 1039 22.3 11.5 49.7 16.6 grad. SH-12 40° 15’ 08” 119° 03’ 47” 1033 21.4 17.6 46.5 14.5 Kbs SH-14 40° 14’ 04” 119° 04’ 27” 1032 25.4 19.2 46.4 9.0 Kbs SH-15 40° 13’ 35” 119° 04’ 48” 906 23.5 22.7 45.3 8.5 Kbs SH-16 40° 13’ 08” 119° 05’ 05” 1086 27.1 17.9 45.6 9.5 Kbs SH-21 40° 12’ 30” 119° 06’ 01” 1090 23.3 20.5 49.7 6.5 Kbs SH-22 40° 12’ 12” 119° 06’ 17” 865 27.9 22.0 41.4 8.8 Kbs SH-23 40° 11’ 49” 119° 06’ 48” 940 26.7 20.5 44.1 8.6 Kbs SH-24 40° 11’ 14” 119° 06’ 46” 1010 22.9 23.0 43.2 11.0 Kbs SH-25 40° 10’ 41” 119° 06’ 14” 1006 26.5 33.6 35.6 4.3 Kbs SH-26 40° 09’ 53” 119° 06’ 48” 1065 24.1 17.0 48.0 10.9 Kbs SH-27 40° 09’ 26” 119° 06’ 24” 903 30.3 24.8 37.4 7.4 Ks SH-29 40° 08’ 26” 119° 05’ 11” 1129 28.3 16.0 44.7 10.9 Ks SH-30 40° 08’ 20” 119° 04’ 33” 827 26.5 16.0 48.1 9.4 Ks NVB-1 40° 07’ 56” 119° 04’ 05” 0 n.d.§ n.d.§ n.d.§ n.d.§

School Bus lobe transect Ksb SB-5 40° 08’ 05” 119° 15’ 43” 877 27.9 24.6 39.5 8.0 Ksb NVB-207 40° 07’ 37” 119° 16’ 06” 977 23.2 23.0 46.3 7.5 Ksb SB-3 40° 07’ 00” 119° 16’ 18” 995 24.0 12.0 53.8 10.3 Ksb SB-2 40° 06’ 25” 119° 16’ 11” 1069 28.4 25.4 40.2 5.9 Ksb SB-1 40° 06’ 02" 119° 16’ 30” 1036 27.1 18.7 48.9 5.2

Aplites aplite AP-02 40° 02’ 45” 119° 08’ 07” 0 n.d.§ n.d.§ n.d.§ n.d.§ aplite AP-03 40° 02’ 47” 119° 08’ 00” 0 n.d.§ n.d.§ n.d.§ n.d.§ aplite AP-04 40° 02’ 50” 119° 07’ 49” 0 n.d.§ n.d.§ n.d.§ n.d.§ aplite AP-05 40° 02’ 26” 119° 03’ 51” 0 n.d.§ n.d.§ n.d.§ n.d.§ aplite AP-06 40° 02’ 28” 119° 03’ 41” 0 n.d.§ n.d.§ n.d.§ n.d.§ aplite AP-07 40° 02’ 27” 119° 03’ 27” 0 n.d.§ n.d.§ n.d.§ n.d.§

†Mineralogy data for sample 10 is not used in Fig. 7 due to insufficient counts. §Not determined. 96

TABLE 3.2. MINERALOGY AND MAJOR ELEMENT CHEMISTRY (CONTINUED)

Unnormalized major elements

SiO2 TiO2 Al2O3 FeO* MnO MgO CaO Na2O K2O P2O5 Total (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %)

Main Sahwave transect 58.74 0.762 18.29 5.33 0.089 2.48 6.01 4.28 1.81 0.265 98.06 60.52 0.839 17.62 5.60 0.095 2.49 5.29 3.95 2.34 0.331 99.07 62.36 0.599 17.34 4.39 0.085 2.00 5.47 4.10 1.95 0.222 98.50 64.16 0.545 16.81 4.09 0.080 1.83 4.66 3.94 2.51 0.195 98.80 65.27 0.548 17.15 3.69 0.070 1.56 4.63 4.18 2.23 0.179 99.53 62.20 0.698 17.19 4.75 0.094 2.10 5.21 3.93 2.45 0.231 98.86 60.51 0.789 18.16 5.30 0.102 2.22 5.51 4.23 2.12 0.262 99.21 60.86 0.791 18.30 4.96 0.093 1.97 5.61 4.19 2.13 0.272 99.18 63.63 0.589 17.79 4.08 0.077 1.51 5.00 4.16 2.05 0.233 99.12 62.66 0.590 17.57 3.90 0.078 1.54 4.52 3.91 3.28 0.212 98.25 65.35 0.442 17.20 2.86 0.057 1.09 3.80 4.38 3.14 0.150 98.48 66.40 0.445 16.65 2.90 0.061 1.05 3.56 4.31 3.09 0.155 98.63 n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ 66.61 0.486 16.31 3.14 0.067 1.10 3.61 4.28 3.06 0.166 98.82 69.10 0.344 15.37 2.25 0.051 0.75 2.85 4.02 3.49 0.123 98.35 n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ 68.55 0.401 15.81 2.40 0.055 0.94 2.92 4.23 3.28 0.130 98.73 n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ 72.42 0.180 14.60 1.26 0.044 0.37 1.79 3.84 4.35 0.060 98.91 n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ 71.14 0.291 14.98 1.92 0.044 0.61 2.59 3.93 3.59 0.113 99.20 67.85 0.303 16.28 1.88 0.039 0.60 2.91 4.07 4.02 0.119 98.08 n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ 68.03 0.333 16.67 2.07 0.049 0.61 2.75 4.27 4.05 0.124 98.94

School Bus lobe transect 70.35 0.226 14.79 1.65 0.052 0.61 2.26 3.39 4.55 0.083 97.95 68.06 0.283 16.10 1.95 0.046 0.68 2.93 3.85 3.96 0.107 97.97 66.68 0.381 16.32 2.58 0.060 0.92 3.45 3.91 3.54 0.137 97.98 n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ n.d.§ 68.30 0.391 15.66 2.49 0.061 0.87 3.01 4.14 3.11 0.126 98.16

Aplites 78.41 0.041 12.21 0.00 0.001 0.04 0.90 2.76 5.17 0.027 99.65 76.84 0.046 13.15 0.13 0.008 0.07 1.12 3.31 4.95 0.025 99.67 77.13 0.052 12.82 0.23 0.003 0.03 0.81 3.34 5.19 0.026 99.65 76.41 0.045 13.10 0.21 0.009 0.02 0.72 3.35 5.79 0.014 99.67 75.47 0.050 13.66 0.33 0.019 0.02 1.15 4.20 4.34 0.016 99.26 77.62 0.055 12.82 0.37 0.010 0.11 1.32 3.15 4.64 0.026 100.19

*All Fe is calculated as FeO. §not determined. 97

Methods

For analysis of modal mineralogy, each sample was sawn into slabs of at least 70 cm2, treated with penetrating epoxy if needed, and ground flat. The slabs were etched with concentrated hydrofluoric acid and stained for K-feldspar and plagioclase using standard procedures as outlined by McMonigle (2002). The stained slabs were sealed with a matte finish, imaged on a flatbed scanner at an effective resolution of 1200 dots per inch, and counted by eye at the intersections of a 2.54 mm grid while enlarged on screen. Areas of significant cracking, epoxy fill, or poor staining were marked off before counting to avoid spurious results. Sample SH-10 and three samples not listed in Table

3.2 were too pervasively cracked to produce enough reliable counting surface. At least

800 points were counted for each of the other samples (Table 3.2). Assuming random counting statistics, this means all 2σ errors should be less than ~5 vol. %, and 2σ errors for modes under 20% should be less than ~3 vol. %.

For geochemical analysis, aliquots of selected samples from the transect were gently hammer crushed before handpicking 30–50 g of fresh chips to be sent to the

Washington State University GeoAnalytical Laboratory for elemental analysis by X-ray fluorescence (XRF) and inductively-coupled-plasma mass spectrometry (Tables 3.2, 3.3).

Additionally, six aplite samples from the southern Sahwave Range were analyzed by

XRF only at the University of California, Santa Cruz.

Three samples were analyzed for Sr and Nd isotopes at the Stanford-USGS Micro

Analysis Center. The samples were prepared by grinding picked chips in a tungsten carbide mill, followed a by HF-HNO3-HCl dissolution procedure in Teflon vials. Sr and

Nd fractions were chemically separated using cation exchange columns in a clean lab TABLE 3.3. TRACE ELEMENT CHEMISTRY

Sample Unnormalized trace elements number Ni Cr V Ga Cu Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm)

Main Sahwave transect SH-1 5 18 130 20 34 92 12.51 29.45 4.15 17.65 3.95 1.13 3.15 0.46 2.56 0.48 1.23 SH-2 8 17 136 22 37 105 13.78 32.53 4.54 18.76 4.11 1.14 3.29 0.47 2.57 0.49 1.26 NVB-206 5 13 109 20 24 81 17.85 34.98 4.34 17.00 3.55 0.98 2.88 0.41 2.30 0.43 1.10 SH-5 4 11 95 19 26 82 8.83 20.75 2.97 12.44 2.72 0.81 2.29 0.33 1.85 0.35 0.93 SH-6 3 10 84 20 13 73 27.68 48.55 5.34 18.95 3.50 0.97 2.74 0.39 2.14 0.41 1.11 SH-7 3 11 112 19 19 89 23.44 49.42 6.32 24.70 5.03 1.26 4.05 0.60 3.27 0.62 1.57 SH-8 5 11 122 21 21 100 25.08 51.19 6.04 22.41 4.32 1.20 3.43 0.50 2.80 0.53 1.40 SH-9 4 9 108 21 8 101 16.66 44.59 6.27 26.10 5.60 1.48 4.53 0.64 3.50 0.65 1.66 SH-10 3 7 83 20 18 81 12.89 33.11 4.72 19.65 4.31 1.17 3.49 0.49 2.76 0.51 1.34 SH-11 b.d.* 7 82 19 13 80 18.49 41.31 5.48 21.54 4.45 1.15 3.46 0.49 2.68 0.50 1.30 SH-12 2 4 60 20 20 67 14.14 30.27 3.84 14.96 3.03 0.85 2.33 0.31 1.63 0.30 0.75 SH-14 1 5 62 19 12 68 16.42 33.19 4.05 15.53 3.01 0.84 2.31 0.31 1.64 0.30 0.78 SH-16 1 4 66 19 14 74 14.11 33.00 4.24 16.53 3.19 0.87 2.45 0.32 1.73 0.32 0.82 SH-21 b.d.* 3 46 17 26 51 20.12 33.05 3.67 13.29 2.52 0.68 1.89 0.25 1.36 0.25 0.65 SH-23 1 5 51 20 10 59 21.64 36.94 4.16 15.24 2.90 0.82 2.21 0.30 1.47 0.26 0.67 SH-25 b.d.* 3 24 18 4 42 14.45 25.18 2.79 9.73 1.79 0.45 1.37 0.19 1.05 0.21 0.58 SH-27 b.d.* 3 38 18 8 48 17.60 29.01 3.19 11.58 2.13 0.58 1.56 0.20 1.03 0.19 0.49 SH-29 b.d.* 3 33 18 8 55 13.96 28.34 3.56 13.63 2.64 0.75 1.89 0.26 1.31 0.24 0.63 NVB-1 1 4 35 20 10 67 15.99 32.99 4.22 16.31 3.13 0.81 2.25 0.28 1.48 0.26 0.63

School Bus lobe transect SB-5 1 4 31 14 1 49 12.60 21.78 2.30 7.78 1.40 0.48 1.12 0.16 0.86 0.17 0.45 NVB-207 b.d.* 4 40 18 6 48 10.93 23.48 2.95 11.25 2.25 0.65 1.71 0.24 1.32 0.26 0.68 SB-3 b.d.* 6 51 18 10 57 15.30 32.58 4.06 15.26 3.02 0.84 2.37 0.33 1.82 0.35 0.94 SB-1 1 5 51 18 8 64 16.09 32.06 3.83 14.35 2.88 0.77 2.12 0.29 1.54 0.28 0.72

Aplites AP-02 n.d.† b.d.* 2 n.d.† n.d.† n.d.† 30 56 n.d.† 20 n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† AP-03 n.d.† 1 3 n.d.† n.d.† n.d.† 30 27 n.d.† 10 n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† AP-04 n.d.† 1 7 n.d.† n.d.† n.d.† 10 19 n.d.† 20 n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† AP-05 n.d.† 1 4 n.d.† n.d.† n.d.† 20 16 n.d.† 10 n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† AP-06 n.d.† 1 4 n.d.† n.d.† n.d.† 40 13 n.d.† 10 n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† AP-07 n.d.† 1 6 n.d.† n.d.† n.d.† 20 44 n.d.† 10 n.d.† n.d.† n.d.† n.d.† n.d.† n.d.† n.d.†

*Below detection. †Not determined. 98 TABLE 3.3. TRACE ELEMENT CHEMISTRY (CONTINUED)

Sample Unnormalized trace elements number Tm Yb Lu Ba Th Nb Y Hf Ta U Pb Rb Cs Sr Sc Zr (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm)

Main Sahwave transect SH-1 0.18 1.09 0.17 1111 1.96 4.67 12.71 3.70 0.27 1.48 9.86 53.6 2.22 808 10.8 139 SH-2 0.18 1.09 0.18 1007 3.71 6.23 12.87 3.21 0.50 1.91 10.77 75.0 2.89 692 9.9 117 NVB-206 0.16 1.00 0.16 876 6.06 5.04 11.43 3.49 0.38 2.06 11.92 55.6 2.51 720 10.4 127 SH-5 0.13 0.85 0.14 819 3.36 4.56 9.29 3.75 0.38 2.45 14.42 83.0 4.15 624 8.7 129 SH-6 0.16 1.04 0.17 803 12.07 6.61 11.15 3.52 0.75 2.80 13.72 72.8 2.85 653 6.8 120 SH-7 0.23 1.49 0.23 1057 7.87 7.73 16.68 3.69 0.95 2.24 13.45 71.5 3.23 672 10.3 132 SH-8 0.20 1.30 0.22 618 10.35 8.38 14.18 4.18 0.87 2.80 11.54 79.5 4.17 663 9.8 153 SH-9 0.25 1.57 0.24 978 5.24 9.64 17.45 3.88 0.84 2.52 12.20 67.8 3.53 800 8.6 144 SH-10 0.19 1.21 0.19 1008 3.99 7.33 13.80 3.58 0.77 1.88 12.23 64.0 3.33 767 6.8 133 SH-11 0.19 1.18 0.19 1947 4.52 7.00 13.80 3.46 0.67 1.66 15.63 81.1 2.85 745 6.4 124 SH-12 0.11 0.74 0.12 1372 4.60 4.93 8.20 3.04 0.48 1.74 19.29 92.2 3.38 647 4.3 107 SH-14 0.11 0.74 0.12 1218 9.23 5.10 8.31 3.12 0.51 3.51 20.51 86.4 2.56 623 4.6 107 SH-16 0.12 0.87 0.15 963 4.31 6.06 9.10 3.54 0.76 2.14 19.99 105.9 8.46 568 5.2 118 SH-21 0.09 0.62 0.11 872 11.88 4.37 7.02 2.88 0.41 3.09 23.33 118.5 6.62 493 3.6 89 SH-23 0.10 0.67 0.11 897 9.21 4.96 7.51 2.94 0.47 3.82 20.85 114.5 6.12 511 3.9 94 SH-25 0.09 0.68 0.13 784 14.41 6.01 6.19 3.48 0.88 4.72 30.65 172.5 5.71 321 2.0 81 SH-27 0.07 0.47 0.08 999 8.69 3.91 5.31 2.58 0.37 2.75 20.28 105.6 3.70 508 2.9 83 SH-29 0.09 0.55 0.09 1573 4.99 4.45 6.67 2.75 0.37 1.59 22.31 104.2 4.11 662 3.1 92 NVB-1 0.09 0.60 0.09 1532 5.96 5.44 7.29 3.33 0.48 1.91 24.52 125.8 7.35 617 3.1 114

School Bus lobe transect SB-5 0.07 0.49 0.09 985 29.34 3.75 4.79 2.62 0.33 5.99 22.39 138.0 5.62 355 3.9 79 NVB-207 0.10 0.69 0.11 1365 4.11 4.49 7.10 2.63 0.48 1.88 22.55 103.4 3.52 544 3.5 88 SB-3 0.14 0.91 0.15 1447 5.19 6.34 9.82 2.79 0.59 1.81 19.17 95.1 4.42 620 4.3 91 SB-1 0.11 0.73 0.12 893 13.41 5.51 7.99 3.20 0.51 3.85 20.70 106.6 5.63 491 3.9 104

Aplites APn.d.†02 n.d.† n.d.† n.d.† 680 n.d.† 3 22 n.d.† n.d.† 3 18 244 n.d.† 306 1 61 APn.d.†03 n.d.† n.d.† n.d.† 230 n.d.† 4 22 n.d.† n.d.† 5 24 225 n.d.† 251 1 63 APn.d.†04 n.d.† n.d.† n.d.† 60 n.d.† 9 23 n.d.† n.d.† 2 33 225 n.d.† 147 2 50 APn.d.†05 n.d.† n.d.† n.d.† 50 n.d.† 6 23 n.d.† n.d.† 6 37 246 n.d.† 105 1 71 APn.d.†06 n.d.† n.d.† n.d.† n.d.† n.d.† 8 19 n.d.† n.d.† 3 31 183 n.d.† 112 3 49 APn.d.†07 n.d.† n.d.† n.d.† 550 n.d.† 4 17 n.d.† n.d.† 1 20 138 n.d.† 339 3 55

†Not determined. 99 100 before loading into a multicollector Finnegan MAT 262 thermal-ionization mass spectrometer on Ta (single) and Re (double) filaments, respectively. Measured 87Sr/86Sr and 143Nd/144Nd ratios were normalized to 86Sr/88Sr = 0.1194 and 146Nd/144Nd = 0.7219 to correct for mass-dependent fractionation. Initial isotope ratios were calculated using

-11 -1 -12 -1 143 144 λ87Rb = 1.42×10 [yr ] and λ147Sm = 6.54×10 [yr ], and Nd/ Ndi is reported as εNd relative to the CHUR evolution model of Jacobsen and Wasserburg (1980).

Mineralogy Results

On a ternary quartz–alkali-feldspar–plagioclase diagram, samples from the

Sahwave Intrusive Suite (circles) define a trend from the quartz diorite field to the granite field, with a majority of the samples falling in the granodiorite field (Fig. 3.7). These mineralogic trends are quite similar to those of the Tuolumne Intrusive Suite (small squares; Fig. 3.7). The Sahwave Intrusive Suite has on average a greater modal abundance of mafic minerals, mainly because mafic granodiorites compose a larger fraction of this intrusion than of the Tuolumne Intrusive Suite. In general, plagioclase and mafic minerals decrease in abundance towards the center of the intrusion while quartz and K-feldspar increase, as expected from field relations (Fig. 3.8). However, these radial modal trends are far from monotonic (Fig. 3.8). For example, the modal percentage of alkali feldspar increases inward for the first 3 km from the contact, drops down to almost its starting value at ~6 km, and then increases to higher values in the

Granodiorite of Bob Spring (Fig. 3.8). The modal percentages of mafic minerals and plagioclase follow a roughly opposite pattern, except plagioclase actually increases inwards from the margin to reach its maximum at ~6 km. Quartz abundance follows 101

Portion of IUGS Classi cation Diagram to Q Sahwave Granodiorite 50 School Bus Granodiorite Granodiorite of Bob Spring granite granodiorite tonalite Granodiorite of Juniper Pass Tuolumne intrusive suite (Bateman, 1992)

quartz quartz quartz monzonite monzodiorite diorite

to 5 A monzonite monzodiorite 35 65 90 P Figure 3.7. Portion of Quartz-Alkali-feldspar-Plagioclase ternary plot showing samples from Sahwave and Tuolumne Intrusive Suites, together with IUGS classifications. 102

Southward distance from intrusion margin (km) 0 5 10 15 20 0 5 70 P School Bus Q Granodiorite of Bob Spring Granodiorite 60 A M

40 Granodiorite of Juniper Pass

30 Mode (vol. %) Mode (vol.

20 Sahwave Granodiorite Sahwave

0 0.0 0.2 0.4 0.6 0.8 1.0 0.0 1.0 % Fraction of radial transect from intrusion margin Figure 3.8. Variations in modal mineralogy as a function of distance from intrusion margin, shown at bottom as a fraction of the total distance; Sahwave Range transect on left, School Bus lobe transect on right. 103 plagioclase’s pattern in reverse. The greatest total variation occurs within the

Granodiorite of Juniper Pass, which is perhaps not surprising given the color index variations and cryptic contacts seen in the field, but variation between individual samples seems to increase in the Granodiorite of Bob Spring (Fig. 3.8). Despite having a greater abundance of K-feldspar megacrysts, the Sahwave Granodiorite is modally quite similar to the Granodiorite of Bob Spring and the School Bus Granodiorite (Fig. 3.8).

Geochemistry Results

Major-element chemistry (Table 3.2) confirms that the Sahwave Intrusive Suite represents a magnesian, metaluminous to weakly peraluminous, calc-alkaline series, with an alkali-lime index of 59.6. Major- and trace-element variation with respect to silica shows trends consistent with fractional crystallization and mixing (open symbols, Fig.

3.9). For example, as differentiation proceeds to higher % SiO2, the incompatible components Rb and K2O increase. Fractional crystallization of hornblende, sphene, and other mafic minerals can explain the decrease in FeO* and Y (not seen in aplite samples, which may have accumulated a Y-rich phase such as xenotime), and plagioclase crystallization keeps Na2O from increasing and removes Sr (Fig. 3.9). Across a radial transect, major elements generally track the same patterns seen in the radial modal plot

(Figs. 3.8, 3.10). On average, major-element chemistry becomes more felsic toward the center of the intrusive complex (Fig. 3.10), but displays significant variation from this general trend. Because these variations fall into a linear array when plotted with respect to silica, much of this local variation may be attributed to mixing between magmas of different composition. Evidence for such mixing is actually observed (Fig. 3.4M) of SiO Figure 3.9.DiagramsshowingvariationofFeO*,K 2008). Na2O (%) K2O (%) FeO* (%) 0 1 2 3 4 5 6 7 0 1 2 3 4 5 6 7 0 1 2 3 4 5 6 7 55 2 inSahwaveand Tuolumne Intrusive Suites(Tuolumne datafromGrayetal., Harker variation diagrams for selected major and trace elements 06 70 65 60 SiO 2 (%) 580 75 Tuolumne intrusive suite aplites Sahwave Granodiorite School BusGranodiorite Granodiorite of Bob Spring Granodirite of Juniper Pass 55 06 70 65 60 2 O, Na 2 SiO O, Rb,Sr, and Y asfunctions 2 (%) 75 80

0 100 200 300 400 500 600 700 800 900 0 50 100 150 200 250 300 0 5 10 15 20 25 30

104

Rb (ppm) Rb Sr (ppm) Sr Y (ppm) Y 105

Sahwave intrusive suite major element transect (with Tuolumne intrusive suite for comparison) 100 SiO2 SiO2

Granodiorite of Juniper Pass Granodiorite of Bob Spring

Al2O3 Al2O3 10

Sahwave Granodiorite Sahwave K2O Na2O K2O Na2O CaO FeO* CaO FeO* 1

log(wt. %) log(wt. MgO

TiO2 MgO

TiO2

P2O5 0.1

P2O5 MnO MnO Half Dome Cathedral Peak Granodiorite Granodiorite Porphyry Johnson Granite 0.01 Quartz Lake of May Diorite 0.0 0.2 0.4 0.6 0.8 1.0 Fraction of radial transect from intrusion margin Figure 3.10. Logarithmic plot showing major-element concentrations as a function of fractional distance along the radial transect from the outside to the center of the intrusion for the Sahwave and Tuolumne Intrusive Suites (Tuolumne data from Bateman and Chappell, 1979). Note that the Tuolumne data are from a shorter, east-west transect of about half the length of the Sahwave transect, but have been expanded as a fraction of radial distance. 106 between diorite intrusions and the Granodiorite of Juniper Pass (which shows the greatest variations) in the southern part of the study area (Fig. 3.3). The Sahwave Intrusive Suite is too long-lived for fractionation of a single large batch of magma after the model of

Bateman and Chappell (1979), but trends towards more felsic major- and trace-element compositions (Fig. 3.9) are consistent with mixing between a set of increasingly fractionated parental magmas.

Compared to the Tuolumne Intrusive Suite, the Sahwave Intrusive Suite tends to have less K2O and Rb, but more Na2O and Sr for any given amount of SiO2 (Fig. 3.9).

Lower K/Na ratios might partially account for the reduced number and size of K-feldspar megacrysts. Rubidium and strontium trends with respect to silica in the Sahwave

Intrusive Suite tend to be somewhat more tightly clustered and show stronger correlations with SiO2 (Fig. 3.9). Compared to the Tuolumne Intrusive Suite, radial major-element variations in the Sahwave Intrusive Suite extend further from the margin and are less monotonic (Fig. 3.10). Compared to the Half Dome Granodiorite, the Granodiorite of

Juniper Pass is almost uniformly richer in Al2O3, CaO, FeO, MgO, and other elements concentrated in plagioclase and mafic minerals (Fig. 3.10). Trends are more similar between the Cathedral Peak and Bob Spring granodiorites, but the Sahwave Intrusive

Suite contains no equivalent to the Johnson Granite Porphyry (Fig. 3.10).

Bulk-rock rare-earth-element (REE) patterns from this transect (Fig. 3.11) all show a broadly similar depression in the heavy REE. The more felsic units are especially depleted in the middle and heavy REE compared to the Granodiorite of Juniper Pass (Fig.

3.11), presumably due to greater fractionation of hornblende, in which these REE are compatible (Arth and Barker, 1976). None of the units show a consistent europium 107

Rare earth element abundance 10,000 Granodiorite of Juniper Pass bulk rock all other Sahwave intrusive suite bulk rock

1000 Granodiorite of Juniper Pass zircon hypothetical liquid in equilibrium with zircons from Granodiorite of Juniper Pass 100

1 Chondrite-normalized abundance

0.1 La Ce NdPr Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Figure 3.11. Chondrite-normalized rare-earth-element diagram comparing whole-rock analyses from the Granodiorite of Juniper Pass with zircons from the same unit and the modeled composition of the liquid that would have been in equilibrium with the zircons. Data from the rest of the Sahwave Intrusive Suite are grouped together for comparison. 108 anomaly. Similar REE patterns in the central Sierra Nevada batholith have been interpreted to reflect differentiation from a deep crustal residue containing garnet (heavy

REE compatible) rather than plagioclase (europium compatible) as the dominant aluminous phase (Ducea, 2001). This hypothesis is also supported by the relatively high

Sr/Y ratio as compared to primitive mantle melts (Fig. 3.12), although arc rocks are generally enriched in fluid-mobile large-ion lithophile elements as compared to relatively high field strength elements (Fig. 3.12). Zircon REE patterns (Fig. 3.11; only

Granodiorite of Juniper Pass zircon data shown for simplicity) show small negative europium anomalies, but this is probably due to the greater incompatibility of europium in zircon due to its 2+ charge, much as the positive cerium anomaly is associated with the

+4 charge taken by those ions. Also shown in Fig. 3.11 is the range of hypothetical liquid compositions that would be in equilibrium with Granodiorite of Juniper Pass zircon, using the partition coefficients for REE in zircon from Sano et al. (2002).

Ignoring lanthanum, the Granodiorite of Juniper Pass REE pattern is within the range of hypothetical liquid compositions, although it is on the upper side, probably because the zircons are frequently included in hornblende, which competes for the middle REE (Fig.

3.11). Zircon saturation temperatures in the range of 748–773 °C were calculated from bulk rock compositions using the model of Watson and Harrison (1983) and provide a likely minimum temperature range for these granodiorite melts.

87 86 Initial Sr/ Sr and εNd (Table 3.4) do not differ greatly among the few measured samples from the Sahwave Intrusive Suite, varying only from 0.7045 to 0.7049 and –

0.256 to –0.173, respectively, indicating a relatively homogeneous source with a strong

87 86 mantle component and little upper crustal assimilation. Sr/ Sr and εNd values are 109

Trace-element abundance 1000 Sahwave intrusive suite

Tuolumne intrusive suite 100

0.1 Large ion N-MORB–normalized abundance lithophile High field strength elements elements

0.01 Sr K BaRb Th Ta Nb P Zr Hf TiSm Y Figure 3.12. Mid-ocean-ridge-basalt-normalized spider diagram showing abundance of trace elements in Sahwave and Tuolumne Intrusive Suites (Tuolumne data from Gray et al., 2008). Large-ion lithophile elements are shown on the left, and high-field-strength elements on the right, with compatibility increasing away from the central dashed line. 110

TABLE 3.4. Sr AND Nd ISOTOPIC DATA

87 86 87 86 143 144 Sample Sr/ Sr ± 2σ Rb Sr Age Sr/ Sri ± 2σ Nd/ Nd ± 2σ Sm Nd εNd ± 2σ number (ppm) (ppm) (Ma) (ppm) (ppm)

SH-1 0.70499±2 53.6 808 92.6 0.70473±2 0.512514±6 3.95 17.65 - 0.173±6 SH-11 0.70528±2 81.1 745 90.5 0.70487±2 0.512466±8 4.45 21.54 - 0.256±8 SH-29 0.70538±2 104.2 662 88.5 0.70480±2 0.512479±9 2.64 13.63 - 0.223±9 S109-6* n.r.† n.r.† n.r.† 92.6 0.7047§ n.r.† n.r.† n.r.† n.r.† S110-6* n.r.† n.r.† n.r.† 88.5 0.7045§ n.r.† n.r.† n.r.† n.r.†

*Sri data for S109-6 and S110-6 (same locations as NVB-206 and NVB-1, respectively) are from Wooden et al. (1999). †Not reported. § Corrections to the reported Sri values, based on the ages from this study and Rb/Sr ratios from co-located samples, are insignificant at the reported precision. 111 similar to those previously measured in the surrounding basinal terrane, which is presumably underlain by relatively mafic transitional crust (e.g., Farmer and DePaolo,

1983), but are considerably more primitive than those measured in similar intrusions along the Sierra Nevada crest (Fig. 3.13).

Analyses of aplite fall near the water-saturated 1 kb haplogranite minimum when projected onto a quartz-orthoclase-albite ternary diagram (Fig. 3.14), consistent with field evidence for the late-stage emplacement of aplite/pegmatite dikes, which likely represent the last fraction of residual melt from a fairly water-rich magmatic system. Scatter between the samples (Fig. 3.14) may be partially due to varying anorthite content and post-magmatic silicification, as well as actual variation in sampled emplacement depth.

When samples containing >5% An or >77.8% SiO2 are excluded, one of the remaining samples still falls significantly below the minimum melt curve, which may suggest the aplite was extracted before water saturation was reached (cf. Nekvasil, 1988), but the other sample (collected 130 m below the Tertiary unconformity) falls just below the 1 kb minimum (Fig. 3.14), consistent with an emplacement depth of 3-4 km. This depth estimate falls within the ca. 3-10 km range suggested by Van Buer et al. (2009) based on the lack of miarolitic cavities and caldera structures and the presence of andalusite in the contact aureole.

A CATHEDRAL RANGE INTRUSIVE EVENT OUTSIDE THE SIERRA

NEVADA

The Sahwave Intrusive Suite is similar to the large intrusions of the Cathedral

Range intrusive epoch in the Sierra Nevada proper (Fig. 3.15) in terms of its ca. 92.5– Nicholas J. Van Buer and Elizabeth L. Miller Figure 13.

112

Sr and Nd isotopic variation Figure 3.13. Strontium and 6 neodymium isotope systematics for 4 in Sahwave and Tuolumne Intru- 2 sive Suites (Tuolumne data from Gray et al., 2008) Also shown are 0

( t ) data from nearby intrusive rocks

Nd -2 within 150 km of the center of the ε -4 Sahwave Intrusive Suite (data from Granodiorite of Juniper Pass Farmer and DePaolo, 1983). -6 Sahwave Granodiorite analyses within 150 km -8 Tuolumne intrusive suite 0.704 0.705 0.706 0.707 Initial 87 Sr/ 86 Sr

Figure 3.14. Quartz-Orthoclase-Albite Q (Q-Or-Ab) ternary diagram comparing measured aplite compositions to experimentally determined minimum-melt compositions (black dots) and phase relations (lines) in a water-saturated system (Johannes and Holtz, 1996), at the pressures labeled in kilobars. Aplite best aplite analyses samples collected from dikes in the aplite analyses with > 5% An or > 77.8% SiO2 Granodiorite of Juniper Pass.

1 2 5

Ab Or

Figure 3.15 (next page): The mid-to-late Cretaceous Sierra Nevada Batholith (darker grey shades) was emplaced across metamorphic rocks and the Jurassic arc (lightest gray). Compositionally zoned intrusions emplaced during the Cathedral Range intrusive epoch (95–83 Ma) are shown as dark grey and black. : D = Domelands Intrusive Suite, JM = John Muir Intrusive Suite, SP = Sonora Pass Intrusive Suite, T = Tuolumne Intrusive Suite, W = Mount Whitney Intrusive Suite, Sahwave Intrusive Suite of this study indicated by S. 113

123º OREGON 122º 121º W 120º W 119º W 118º W CALIFORNIA OREGON 42º N 2004 Stanford seismic line NEVADA (Lerch et al., 2007) colored by inferred basement

KLAMATH RANGE 41º N

Selenite R. area of Fig. 3 S Trinity R. 40º N

Reno Lake B A S I N Tahoe A N D S I E R R A N E V A D A R A N G E 39º N Sacramento

G R E A T V A L L E Y SP ? CALIFORNIANEVADA

T 38º N

Yosemite Village

37º N

W Explanation of Map Units

Cenozoic cover 36º N Younger zoned plutons, 94-83 Ma

Older zoned plutons, 94-83 Ma D

Unexposed zoned plutons?

Other Mesozoic intrusions

Other pre-Cenozoic rocks

Geologic map compiled mainly from Jennings et Garlock fault al., 1977; Irwin and Wooden, 2001; Bateman, San Andreas fault 50 km 1992; John, 1983; Tikoff and de Saint Blanquat, 1997; Hirt, 2007; Saleeby et al., 2008. Ages are from many sources. 114

88.5 Ma age range, >1000 km2 size, modal and chemical zonation, and internal magmatic structures. It is different in that it has a somewhat lower K/Na ratio, smaller abundance and lesser size of K-feldspar megacrysts, a larger proportion of mafic granodiorite, more primitive Sr and Nd isotopic ratios, its relatively equant rather than elongate shape, and its location to the north and east of the Sierra Nevada crest (Fig. 3.15). Additionally, the approximately four-million-year apparent duration of the Sahwave Intrusive Suite is shorter duration than that reported for many of the coeval intrusive suites along the Sierra

Nevada crest, but it is similar to the durations represented by the most voluminous phases of those suites (Chen and Moore, 1982; Coleman et al., 2004; Saleeby et al., 2008).

Many of these intrusive suites, e.g., the Tuolumne Intrusive Suite, had shorter reported durations of intrusion before extensive geochronology campaigns were carried out (e.g.,

Coleman et al., 2004), so it is possible that the Sahwave Suite may also include minor phases that would extend its reported duration. Despite some differences, the similarities are compelling enough to consider the Sahwave Intrusive Suite a member of the

Catherdal Range supersuite. The continuation of this distinctive chain of intrusions into the NW Basin and Range further supports the idea of an originally continuous Sierra

Nevada Batholith later disrupted by Cenozoic extension (Fig. 3.15). The wide separation between the Sahwave Intrusive Suite and the Sonora Pass Intrusive Suite shown in Fig.

3.15 may suggest the existence of another intrusive suite (or suites) in the Reno area, where little-studied granitoids of similar age are exposed (91–86 Ma K/Ar biotite ages;

Marvin and Cole, 1978; Garside et al., 1992).

115

Emplacement

The emplacement mechanisms of these large, long-lived intrusions remain controversial. In light of detailed geochronology, geochemistry, and numerical modeling, a variety of subtle internal structures have been generally interpreted to suggest that these intrusive complexes are formed by repeated influx of magma batches into a system kept near its solidus (e.g., Coleman et al., 2004; Hirt, 2007; Saleeby, 2008). However, opinions differ greatly on the size, frequency, and emplacement mechanisms of these magmatic replenishment events, varying from frequent re-intrusion of small dikes (e.g.

Glazner et al., 2004) to much larger batches of magma that remain above their solidi for extended periods of time (e.g., Zak and Paterson, 2005). We briefly evaluate evidence for different emplacement mechanisms in the Sahwave Intrusive Suite.

The Sahwave Intrusive Suite generally has steep contacts and steep magmatic foliation (Fig. 3.3); thus it is unlikely that it is a sill-like intrusion. Nevertheless, given its

40 km diameter at a relatively shallow depth of exposure, the intrusion must have been relatively flat-topped (Fig. 3.16). The downward extent of the batholith is not well defined by existing seismic or gravity data, but comparison to the oblique crustal arc sections exposed in southern California suggests that batholithic rocks may extend to the base of the crust, although large distinct intrusions in the upper crust may overlie a complex zone of smaller, vertically sheeted intrusions in the lower crust (Saleeby, 2003;

Barth et al., 2008; Saleeby et al., 2008). Seismic data from further north along the arc

(Fig. 3.2; Lerch et al., 2007), however, indicates low velocities compatible with tonalitic/granitic rocks down to ~15 km, suggesting that the magmatic arc in NW Nevada is underlain by mafic residua and remnants of thin transitional crust, in contrast to the indicates presentleveloferosion. remaining mushyandhomogeneousbecausethe systemwarmedovertime.Dashedline model ofSaleeby(2008). The later, centralplutons wereabletoreachlarger sizes while mixture ofsmaller, verticallyalignedintrusiveremnantsinthelowercrust,following Cretaceous. A fewlarge, nestedplutonsintheuppercrust arehypothesizedtooverliea Figure 3.16.CartooncrosssectionofSahwaveIntrusiveSuite afterintrusionintheLate

Depth (km) 40 30 20 10 0 5 km of JuniperPass Granodiorite basaltic underplating and partial basaltic underplating andpartial remelting (minorassimilation), Granodiorite storage, andhybridization? Crustal-scale cartoon thick root ofma cresidua Sahwave of Bob Spring of Bob Granodiorite 116 117 crust of the southern and central Sierra Nevada Batholith, which is relatively felsic

(tonalitic) to its base (e.g., Saleeby, 1990; Fleidner, 2000). Even if the batholithic rocks studied here extend to only 15 km depth, as shown in Fig. 3.16, The Sahwave Intrusive

Suite still represents a volume of well over 10,000 km3, such that space accommodation and mechanisms of its emplacement are non-trivial problems.

Given the vast size of the Granodiorite of Juniper Pass and the variations in modal mineralogy and chemistry it contains, it is entirely possible that this unit was emplaced over time as a series of smaller intrusive events. The cryptic internal structures and contacts within the Granodiorite of Juniper Pass are difficult to interpret, but the general smoothness of compositional variation within the sampled part of this unit suggests that individual batches of magma generally stayed hot long enough to partially mix with their successors. Other contacts within and around the Sahwave Intrusive Suite also vary greatly in style. Although often poorly exposed, contacts between units can be both sharp, such as where the School Bus Granodiorite intrudes the Granodiorite of Juniper

Pass, or gradational over hundreds of meters, such as where the Granodiorite of Bob

Spring intrudes the Granodiorite of Juniper Pass (Fig. 3.3). In certain places, contacts are observed to transition from sharp to gradational. This happens gradually along the contact of the Sahwave Granodiorite, but fairly abruptly where the Granodiorite of Bob Spring intrudes the Granodiorite of Juniper Pass east of the Power Line Complex in the

Nightingale Range (Fig. 3.3; Plates 4, 8).

Sharp internal contacts likely reflect areas where new magma batches “eroded” their way into older magma that had partially cooled to the point where it behaved as a solid, perhaps even experiencing brittle fracture and stoping in places. Conversely, 118 arcuate, gradational contacts likely formed where the previous batch of magma was either still partially molten, or at least close enough to its solidus, to experience defrosting and partial to complete mixing along its contact. The apparent lack of internal contacts within the Granodiorite of Bob Spring, the School Bus Granodiorite, and the Sahwave

Granodiorite (with the exception of a few contacts surrounding leucocratic segregations), combined with the general homogeneity of these units, suggests that each may represent a single phase of rapid magma input into a large, partially molten magma chamber. The concentric arrangement of these units further suggests that the central part of each unit was not fully mechanically and/or thermally stabilized before its successor intruded, and possibly flowed back downwards after defrosting to accommodate the new magma. The concentric arrangement of successively more homogeneous (and generally more differentiated) units (Fig. 3.16) also suggests that the system was warming over time, allowing larger and longer-lived magma chambers to be formed at both the level of exposure and perhaps at the deeper level of magma production. Warming over time in the southern Nightingale Range due to reintrusion of the School Bus Granodiorite so close to the contact already heated once by the Granodiorite of Juniper Pass may also be responsible for the extensive shouldering-aside implied by the anomalous orientation of the wall rocks in this area (Fig. 3.3).

Basinal Setting and Implications for Arc Flare-Up

In light of the extensive similarities between the Sahwave Intrusive Suite and coeval magmatic systems to the south, it is interesting to note that the crustal environments of these intrusions are very different (Figs. 3.1, 3.15). Whereas the other 119 massive intrusions of the Cathedral Range intrusive epoch are interpreted to lie along the margin of North American continental crust, as marked by scattered roof pendants of the miogeocline and initial 87Sr/86Sr > 0.706 (Fig. 3.1; e.g., Saleeby, 1981; Kistler, 1990), the

Sahwave Intrusive Suite is positioned in a deep stack of basinal muds thought to overlie transitional or oceanic crust (Speed, 1978; Farmer and DePaolo, 1983; Elison et al.,

1990). Because of its unique position relative to the other members of the Cathedral

Range event, the Sahwave Intrusive Suite can be used to examine hypotheses about the potential causes for the massive magmatic flare-up represented by these intrusions. It has been suggested that this particular pulse of major magmatic activity may be due to westward underthrusting of North American lower crust beneath the magmatic arc, which is hypothesized to have been near the western edge of a massive orogenic wedge

(DeCelles and Coogan, 2006; DeCelles et al., 2009). The voluminous magmatism in the

Sahwave and Nightingale Ranges at this time, however, demonstrates very primitive

87 86 isotopic ratios of Sr/ Sri ~ 0.7047 and εNd ~ -0.2, which are not compatible with incorporation of a large crustal component. Similar or more primitive isotope ratios in the penecontemporaneous La Posta events of the Peninsular Ranges Batholith (Fig. 3.1;

Walawender et al., 1990) corroborate this. Furthermore, modest reconstructed

Cretaceous crustal thicknesses near the Sahwave batholith (~38 km; e.g., Colgan et al.,

2006) would suggest that the orogenic wedge did not continue this far west in northern

Nevada.

Since the availability of continental lower crust does not appear to be the main control on high magmatic flux in the arc at this time, it seems that a more regionally extensive and consistent triggering mechanism must be invoked. A widespread flare-up 120 in the arc could be related to the tectonic underplating of Franciscan subduction accretionary material (Saleeby et al., 2008), a change in subduction rate and/or obliquity, age or composition of underthrust oceanic lithosphere, or the stress regime accompanying intrusion. Alternatively, subduction of thicker oceanic crust in the Late Cretaceous could potentially be called upon to both induce a magmatic flux event and subsequently terminate magmatism. Very shallow subduction of a large and thick oceanic plateau is hypothesized to have disrupted the Mojave-Salinia segment of the arc (Saleeby, 2003), but modestly thickened oceanic crust in adjacent segments might have lead to moderately shallow subduction and the observed cessation of magmatism. Thicker oceanic crust might incorporate and react with a greater volume of seawater (especially if pillow basalts represent a disproportionate share of crustal thickening in oceanic plateaus; e.g., Gladczenko et al., 1997). Dragged downwards by previously-subducted, denser oceanic lithosphere, the leading edge of this thickened oceanic crust could release its fluids into the mantle wedge, creating a massive, fluid-rich basaltic flux that might remelt any stack of older basalt left underplated at the base of the arc crust by prior arc activity.

Triggered by the same cause as the incipient flat-slab subduction, magmatic flux would increase until crowding from the increasingly buoyant oceanic lithosphere caused stagnation of the mantle wedge, halting magmatism.

CONCLUSIONS

New mapping, geochronology, petrology and geochemistry in the Sahwave and

Nightingale Ranges of western Nevada document the northward continuation of the

Cretaceous Cordilleran arc across the NW Basin and Range and lay the groundwork for 121 more detailed future study. Intrusive activity in the Sahwave and Nightingale area continued from ca. 110–88.5 Ma, and included the emplacement of a large, concentrically zoned intrusive suite at ca. 93-88.5 Ma, during the culminating magmatic flare-up of the

Sierra Nevada Batholith. Mineralogy and geochemistry support the correlation of the

Sahwave Intrusive Suite with members of the Cathedral Range intrusive event along the

Sierra Nevada crest. The oldest unit of the Sahwave Intrusive Suite, the Granodiorite of

Juniper Pass, is marked by significant compositional variations, which may indicate formation from multiple smaller intrusions, but the later, K-feldspar-porphyritic

Granodiorite of Bob Spring and Sahwave Granodiorite are more homogeneous and may represent relatively large magma-mush chambers that were continuously maintained above their solidi. Concentric arrangement and gradational contacts between different units imply that parts of the system remained near the solidus during much of the four million years it was active. Despite differences in the lower crust beneath the Sahwave

Intrusive Suite and intrusions along the crest of the Sierra Nevada (as indicated by more

87 86 primitive Sr/ Sri and εNd values to the north), striking similarities between these segments of the arc suggest that a regionally developed sub-crustal mechanism, such as the subduction of thicker and wetter oceanic crust or Franciscan mélange, may be responsible for generation of the intrusions which punctuate the end of Cretaceous magmatism along much of the U.S. Cordilleran arc.

ACKNOWLEDGEMENTS

This research was partially sponsored by NSF Tectonics grant 0809226, two Stanford

McGee Grants, and a GSA Student Research Grant. Van Buer was partially supported by 122 a Burt and DeeDee McMurtry Fellowship. Special thanks to Joe Wooden for help acquiring and analyzing SHRIMP data, to Bettina Wiegand for measuring Sr and Nd isotopes, and to other helpful folks at the Stanford-U.S.G.S. Micro-Analysis Facility.

Thanks also for helpful reviews from Gail Mahood, Robinson Cecil, and Sandra Wyld.

REFERENCES

Arth, J. G., Barker, F., 1976, Rare-earth partitioning between hornblende and dacitic liquid and implications for the genesis of trondhjeimitic-tonalitic magmas, Geology, v. 4, p. 534-536.

Barth, A.P., Anderson, J.L., Jacobsen, C.E., Paterson, S.R., Wooden, J.L., 2008, Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert: Geological Society of America Field Guide, v. 11, p. 101-117.

Bateman, P.C., 1992, Plutonism in the central part of the Sierra Nevada Batholith, California: U.S. Geological Survey Professional Paper 1483.

Bateman, P.C., Chappell, B.W., 1979, Crystallization, fractionation, and solidification of the Tuolumne intrusive series, Yosemite National Park, California: Geological Society of America Bulletin, v. 90, p. 465- 482.

Black, L.P., Kamo, S.L., Allen, C.M., Davis, D.W., Aleinikoff, J.N., Valley, J.W., Mundil, R., Campbell, I.H., Korsch, R.J., Williams, I.S., and Foudolis, C., 2004, Improved 206Pb/238U microprobe geochronology by the monitoring of a trace- element-related matrix effect: SHRIMP, ID-TIMS, ELA-ICP-MS, and oxygen isotope documentation for a series of zircon standards: Chemical Geology, v. 205, p. 115-140.

Bonham, H.F., 1969, Geology and Mineral Deposits of Washoe and Storey Counties, Nevada: Nevada Bureau of Mines Bulletin, v. 70, 140 pp.

Burke, D.B., and Silberling, N.J., 1973, The Auld Lang Syne Group, of Late Triassic and 123

Jurassic(?) age, north-central Nevada: Contributions to Stratigraphy, United States Geological Survey Bulletin 1394-E.

Busby-Spera, C., and Saleeby, J.B., 1990, Intra-arc strike-slip fault exposed at batholithic levels in the southern Sierra Nevada, California: Geology, v. 18,p. 255-259.

Chen, J.H., Moore, J.G., 1982, Uranium-lead isotopic ages from the Sierra Nevada Batholith, California: Journal of Geophysical Research, v. 87, p. 4761-4784.

Coleman, D.S., and Glazner, A.E., 1998, The Sierra Crest magmatic event: Rapid formation of juvenile crust during the Late Cretaceous in California, in Hall, C.A., Ernst, W.G., and Nelson, C.A., eds., Integrated Earth and Environmental Evolution of the Southwestern United States: The Clarence A. Hall, Jr., Volume: The Geological Society of America, p. 253-272.

Coleman, D.S., W. Gray, A.F. Glazner, 2004, Rethinking the emplacement and evolution of zoned plutons; geochronologic evidence for incremental assembly of the Tuolumne Intrusive Suite, California, Geology, v. 32 (5), p. 433-436.

Colgan, J.P., Wyld, S.J., and Wright, J.E., 2010, Geologic map of the Vicksburg Canyon quadrangle, Humboldt County, Nevada: Nevada Bureau of Mines and Geology Map 169.

Compton, R.R., 1960, Contact metamorphism in Santa Rosa Range, Nevada: Geological Society of America Bulletin, v. 71, p. 1383-1416.

DeCelles, P.G., and Coogan, J.C., 2006, Regional structure and kinematic history of the Sevier fold-and-thrust belt, central Utah: Geological Society of America Bulletin, v. 118, p. 841-864.

DeCelles, P.G., Ducea, M.N., Kapp, P., Zandt, G., 2009, Cyclicity in Cordilleran orogenic systems: Nature Geoscience, v. 2, p. 251-257.

DeGraaff-Surpless, K., Graham, S.A., Wooden, J.L., McWilliams, M.O., 2002, Detrital zircon provenance analysis of the Great Valley Group, California: Evolution on 124

an arc-forearc system: Geological Society of America Bulletin, v. 114, p. 1564- 1580. du Bray, E.A., 2007, Time, space, and composition relations among northern Nevada intrusive rocks and their metallogenic implications: Geosphere, v. 3 (5), p. 381- 405.

Ducea, M., 2001, The California arc: Thick granitic batholiths, eclogitic residues, lithospheric-scale thrusting, and magmatic flare-ups: GSA Today, v. 11, p. 4-10.

East, J.H., and Trengrove, R.R., 1950, Investigation of the Nightingale tungsten deposit, Pershing County, Nevada: U.S. Bureau of Mines report of investigations 4678.

Elison, M.W., Speed, R.C., Kistler, R.W., 1990, Geologic and isotopic constraints on the crustal structure of the northern Great Basin: GSA Bulletin, v. 102, p. 1077-1092.

Engebretson, D.C., Cox, A., Gordon R.G., 1985, Relative motions between oceanic and continental plates in the Pacific Basin: Geological Society of America Special Paper 206.

Evernden, J.F., and Kistler, R.W., 1970, Chronology of emplacement of Mesozoic batholith complexes in California and western Nevada: U.S. Geological Survey Professional Paper 623.

Fanning, D.J., 1982, Metamorphism and tungsten mineralization in the Nightingale Range, Pershing County, Nevada: [thesis; U Nevada Reno].

Faulds, J.E., Henry, C.D., Hinz, N.H., 2005, Kinematics of the northern Walker Lane; an incipient transform fault along the Pacific-North American Plate boundary: Geology, v. 33, p. 505-508.

Fliedner, M., S. Klemperer, and N. Christensen (2000), Three-dimensional seismic model of the Sierra Nevada Arc, California, and its implications for crustal and upper mantle composition, J. Geophys. Res., 105, 10,899 – 10,921.

Fowler, D.D., Fowler C.S., 1971, Anthropology of the Numa: John Wesley Powell’s Manuscripts on the Numic Peoples of Western North America 1868-1880. Washington, D.C.: Smithsonian Institution Press.

125

Garside, L.J., Bonham, H.F., and McKee, E.H., 1992, Potassium-argon ages of plutonic rocks and associated vein and alteration minerals, northeast Sierra Nevada, California and Nevada: Isochron/West, v. 58, p. 13-15.

Gastil, R.G., 1975, Plutonic zones in the Peninsular Ranges of southern California and northern Baja California: Geology, v. 3 (7), p. 361-363.

Gastil, G., 1983, Mesozoic and Cenozoic granitic rocks of southern California and western Mexico: Memoir – Geological Society of America, v. 159, p. 265-275.

Gladczenko, T.D., Coffin, M.F., Eldholm O., 1997, Crustal structure of the Ongtong Java Plateau: Modeling of new gravity and existing seismic data: Journal of Geophysical Research, v. 102, p. 22,711-22-729.

Glazner, A.F., Bartley, J.M., Coleman, D.S., Gray, W., and Taylor, R.Z., 2004, Are plutons assembled over millions of years by amalgamation from small magma chambers: GSA Today, v. 14 (5), p. 4-11.

Gray, W., Glazner, A.F., Coleman, D.S., and Bartley, J.M., 2008, Long-term geochemical variability of the Late Cretaceous Tuolumne Intrusive Suite, central Sierra Nevada, California, in Annen, C., Zellmer, G.F., eds., Dynamics of Crustal Magma Transfer, Storage, and Differentiation: Geological Society, London, Special Publication 304, p. 183-201.

Greene, D.C., Schweikert, R.A., 1995, The Gem Lake shear zone: Cretaceous dextral transpression in the Northern Ritter Range pendant, eastern Sierra Nevada, California: Tectonics, v. 14, p. 945-961.

Hildreth, W., and Moorbath, S., 1988, Crustal contributions to arc magmatism in the Andes of central Chile: Contributions to Mineralogy and Petrology, v. 98, p. 455- 489.

Hirt, W.H., 2007, Petrology of the Mount Whitney Intrusive Suite, eastern Sierra Nevada, California; implications for the emplacement and differentiation of composite felsic intrusions: GSA Bulletin, v. 119 (9-10), p. 1185-1200.

Hyndman, D.W., 1983, The Idaho Batholith and associated plutons, Idaho and western Montana, in Roddick, J.A., ed., Circum-Pacific plutonic terranes: Geological Society of America Memoir 159, p. 213-140.

126

Jacobsen, S.B., Wasserburg, G.J., 1980, Sm-Nd isotopic evolution of chondrites: Earth and Planetary Science Letters, v. 50, p. 139-155.

Jennings, C., R. Strand, and T. Rogers, 1977, Geologic map of California, scale 1:750,000, Calif. Div. of Mines and Geol., Sacramento.

Johannes, W., Holtz, F., 1996, Petrogenesis and Experimental Petrology of Granitic Rocks. Berlin: Springer, 335 pp.

John, D.A., 1983, Distribution, ages, and petrographic characteristics of Mesozoic plutonic rocks, Walker Lake 1º by 2º quadrangle, California and Nevada: Miscellaneous Field Studies, Map MF-1382-B.

John, D.A., 1992, Chemical analyses of granitic rocks in the Reno 1º by 2º quadrangle and in the northern Pine Nut mountains, west-central Nevada, USGS Open-File Report 92-246.

John, D.A., Robinson, A.C., 1982, Comagmatic granitoid sequence, Sierra Nevada: U. S. Geological Survey Professional Paper, Report P1375, p. 67.

John, D.A., Schweikert, R.A., and Robinson, A.C., 1994, Granitic Rocks in the Triassic- Jurassic Magmatic Arc of Western Nevada and Eastern California: U.S. Geological Survey Open-File Report 94-148, 61 pp.

Johnson, M., 1977, Geology and mineral deposits of Pershing County, Nevada: Nevada Bureau of Mines and Geology Bulletin 89.

King, P.B., Beikman, H.M., 1974, Geologic map of the United States (exclusive of Alaska and Hawaii) on a scale of 1:2,500,000: U.S. Geological Survey.

Kistler, R.W., 1990, Two different lithosphere types in the Sierra Nevada, California, in Anderson, J.L., ed., The nature and origin of Cordilleran magmatism: Boulder, Colorado, GSA Memoir 174, p. 271-281.

Kistler, R.W., 1999, Plutons of the Late Cretaceous Cathedral Range Intrusive Epoch: Testing ground for models of the origin of the composite Sierra Nevada Batholith, California: Abstracts with Programs – Geological Society of America, v. 31 (6), p. 70.

Kistler, R.W., Chappell, B.W., Peck, D.L., Bateman, P.C., 1986, Isotopic variation in the Tuolumne Intrusive Suite, central Sierra Nevada, California: Contributions in Mineralogy and Petrology, v. 94, p. 205-220.

127

Lackey, J.S., Valley, J.W., Saleeby, J.B., 2005, Supracrustal input to magmas in the deep crust of Sierra Nevada Batholith; evidence from high δ18O zircon: Earth and Planetary Science Letters, v. 235, p. 315-330.

Ludwig, K.R., 2001, Squid: A users’ manual: Berkeley Geochronology Center Special Publication 2, 19 p.

Ludwig, K.R., 2003, Isoplot 3.00, a geochronological tool-kit for Excel: Berkeley Geochronology Center Special Publication 4, 67 pp.

Marvin, R.F., Cole, J.C., 1978, Radiometric ages: Compilation A, U.S. Geologic Survey: Isochron/West, v. 22, p. 3-14.

McMonigle, M., Miller, K., Monroe, M., 2002, Feldspar staining of the plutonic rocks of the Salinian Terrane: http://www.marin.edu/~jim/geolprojects/ (May 2008).

Morton, J.L., Silberman, M.L., Bonham, H.F., Jr., Garside, L.J., and Noble, D.C., 1977, K-Ar ages of volcanic rocks, plutonic rocks, and ore deposits in Nevada and eastern California; determinations run under the USGS-NBMG Cooperative Program: Isochron/West, v. 20, p. 19-29.

Nekvasil, H., 1988, Calculated effect of anorthite component on the crystallization paths of H2O-undersaturated haplogranitic melts, American Mineralogist, v. 73, p. 966- 981.

Oldenburg, E.A., 1995, Chemical and Petrologic Comparison of Cretaceous Plutonic Rocks from the Diamond Mountains, Fort Sage Mountains and Yuba Pass Region of Northeastern California: Humboldt State University [MS thesis].

Oldow, J.S., 1984, Evolution of a late Mesozoic back-arc fold and thrust belt, northwestern Great Basin, U.S.A.: Tectonophysics, v. 102, p. 245-274.

128

Rai, V.N., 1969, Geology of a portion of the Nightingale and Truckee ranges, Washoe and Pershing counties, Nevada: [thesis].

Saleeby, J.B., 1981, Ocean floor accretion and volcanoplutonic arc evolution of the Mesozoic Sierra Nevada, California, in Ernst, W.G., ed., Rubey Volume on the Geotectonic Development of California: Englewood Cliffs, New Jersey, Prentice-Hall, p. 132-181.

Saleeby, J.B., 1990, Progress in tectonic and petrogenetic studies in an exposed cross- section of young (~100 Ma) continental crust, southern Sierra Nevada, California, in Salisbury, M.H., ed., Exposed Cross Sections of the Continental Crust: Dordrecht, Netherlands, D. Riedel Publishing, p. 137-158.

Saleeby, J.B., 2003, Segmentation of the Laramide Slab—evidence from the southern Sierra Nevada region: Geological Society of America Bulletin, v. 115, p. 655- 668.

Saleeby, J.B., Ducea, M.N., Clemens-Knott D., 2003, Production and loss of high-density batholithic root, southern Sierra Nevada, California: Tectonics, v. 22, doi: 10.1029/2002TC001374.

Saleeby, J.B., Ducea, M.N., Busby, C.J., Nadin, E.S., Wetmore, P.H., 2008, Chronology of pluton emplacement and regional deformation in the southern Sierra Nevada batholith, California, in Wright, J.E., Shervais, J.W., eds., Ophiolites, Arcs, and Batholiths: A Tribute to Clif Hopson: Geological Society of America Special Paper 438, p. 397-427.

Sano, Y., Terada, K., Fukoka, T., 2002, High mass resolution ion microprobe analysis of rare earth elements in silicate glass, apatite and zircon; lack of matrix dependency: Chemical Geology, v. 184, p. 217-230.

Silver, L.T., Taylor Jr., H.P., Chappell, B.W., 1979, Some petrological, geochemical, and geochronological observations of the Peninsular Ranges batholith near the international border of the U.S.A. and Mexico, in Abbott, P.L., and Todd, V.R., eds., Mesozoic Crystalline Rocks: Peninsular Ranges Batholith and Pegmatites, Point Sal Ophiolite: Geological Society of America Guidebook, p. 83-110.

Smith, J.G., McKee, E.H., Tatlock, D.B., Marvin, R.F., 1971, Mesozoic granitic rocks in 129

northwestern Nevada: A link between the Sierra Nevada and Idaho Batholiths: GSA Bulletin, v. 82, p. 2933-2944.

Smith, W.C., and Guild, P.W., 1942, Tungsten deposits of the Nightingale district, Pershing County, Nevada: U.S. Geological Survey Bulletin 936-B.

Stager, H.K., and Tingley, J.V., 1988, Tungsten deposits of Nevada: Nevada Bureau of Mines and Geology Bulletin 105.

Streckeisen, A.L., 1976, To each plutonic rock its proper name: Earth Science Reviews, v. 12, p. 1-33.

Tikoff, B., de Saint Blanquat, M., 1997, Transpressional shearing and strike-slip partitioning in the Late Cretaceous Sierra Nevada magmatic arc, California: Tectonics, v. 3, p. 442-459.

Tikoff, B. and C. Teyssier, 1992, Crustal-scale, en echelon “P-shear” tensional bridges; a possible solution to the batholithic room problem, Geology, v. 20 (10), p. 927-930.

Titus, S.J., R. Clark, B. Tikoff, 2005, Geological and geophysical investigation of two fine-grained granites, Sierra Nevada Batholith, California; evidence for structural controls on emplacement and volcanism, GSA Bulletin, v. 117 (9-10), p. 1256- 1271.

Van Buer, N.J., and Wooden, J.L., 2007, An Upper Cretaceous compositionally zoned intrusion in the Sahwave and Nightingale Ranges, Pershing County, Nevada: Geological Society of America – Abstracts with Programs, v. 39 (4), p. 22.

Van Buer, N.J., Miller, E.L., Dumitru, T.A., 2009, Early Tertiary paleogeologic map of the northern Sierra Nevada batholith and the northwestern Basin and Range: Geology, v. 37, p. 371-374.

Walawender, M.J., Gastil, R.G., Clinkenbeard, J.P., McCormick, W.V., Eastman, B.G., Wernicke, R.S., Wardlaw, M.S., Gunn, S.H., and Smith, B.M., 1990, Origin and evolution of the zoned La Posta-type plutons, eastern Peninsular Ranges batholith, southern and Baja California, in Anderson, J.L., ed., The nature and origin of Cordilleran magmatism: Boulder, Colorado, Geological Society of America Memoir 174.

130

Watson, E.B., Harrison, T.M., 1983, Zircon saturation revisited; temperature and composition effects in a variety of crustal magma types: Earth and Planetary Science Letters, v. 64, p. 295-304. Wesnousky, S.G., 2005, The San Andreas and Walker Lane fault systems, western North America: transpression, transtension, cumulative slip and the structural evolution of a major transform plate boundary: Journal of Structural Geology, v. 27, p. 1505- 1512.

Whitehill, C. S., 2009, Cenozoic evolution of the Shawave-Nightingale horst block, northwestern Basin and Range, Nevada, U.S.A. [thesis].

Willden, R., 1964, Geology and Mineral Deposits of Humboldt County, Nevada: Nevada Bureau of Mines and Geology Bulletin, v. 59, 154 pp.

Willden, R., and Speed, R.C., 1974, Geology and Mineral Deposits of Churchill County, Nevada, v. 83, 95 pp.

Wyld, S.J., 1996, Early Jurassic deformation of the Pine Forest Range, northwest Nevada, and implications for Cordilleran tectonics: Tectonics, v. 15, p. 566-583.

Wyld, S.J., 2000, Triassic evolution of the arc and backarc of northwest Nevada, and evidence for extensional tectonism, in Soreghan, M.J., and Gehrels, G.E., eds., Paleozoic and Triassic Paleogeography and Tectonic Evolution of Western Nevada and Northern California: Boulder, Colorado, Geological Society of America Special Paper 347, p. 185-208.

Wyld, S.J., Rogers, J.W., Copeland, P., 2003, Metamorphic evolution of the Luning- Fencemaker Fold-Thrust Belt, Nevada: Illite crystallinity, metamorphic petrology, and 40Ar/39Ar geochronology: The Journal of Geology, v. 111, p. 17-38.

Wyld, S.J., and Wright, J.E., 1997, Triassic-Jurassic tectonism and magmatism in the Mesozoic continental arc of Nevada: Classic relations and new developments, in Line, P.K., and Kowallis, B.J., eds., Proterozoic to Recent Stratigraphy, Tectonics, and Volcanology in Utah, Nevada, Southern Idaho, and Central Mexico: Geological Society of America Field Trip Guide Book, Brigham Young University 131

Geology Studies, v. 42 (1), p. 197-224.

Zak, J., and S.R. Paterson, 2005, Characteristics of internal contacts in the Tuolumne Batholith, central Sierra Nevada, California (USA); implications for episodic emplacement and physical processes in a continental arc magma chamber, GSA Bulletin, v. 117 (9-10), p. 1242-1255.

132

APPENDIX A: U-Pb SHRIMP ANALYTICAL DATA

TABLE A1. U-Pb SHRIMP ANALYTICAL DATA

Spot number Common 206Pb U Th 232Th/238U 206Pb/238U Age ± 1σ Total 238U/206Pb (± Total 207Pb/206Pb (± (%) (ppm) (ppm) (Ma) % err) % err)

JC03-SV03: Sahwave Granodiorite SV3-1 0.12 623 286 0.47 84.8 ± 1.3 75.4 ± 1.5 .0486 ± 3.3 SV3-2 0.08 667 241 0.37 87.3 ± 1.3 73.3 ± 1.5 .0484 ± 3.2 SV3-3 0.28 497 106 0.22 86.3 ± 1.4 74.0 ± 1.6 .0499 ± 3.8 SV3-4 0.13 750 245 0.34 87.7 ± 1.3 72.9 ± 1.5 .0488 ± 3.1 SV3-5 0.11 669 199 0.31 87.9 ± 1.3 72.8 ± 1.5 .0486 ± 3.2 SV3-6 0.37 601 180 0.31 90.1 ± 1.4 70.8 ± 1.6 .0508 ± 3.3 SV3-7 0.11 867 348 0.42 91.4 ± 1.3 70.0 ± 1.5 .0487 ± 2.8 SV3-8 -0.40 753 244 0.33 91.6 ± 1.4 70.1 ± 1.5 .0447 ± 3.1 SV3-9 0.25 529 161 0.31 89.9 ± 1.4 71.0 ± 1.6 .0498 ± 3.4

NVB-207: School Bus Granodiorite NVB207-7 -0.10 2081 457 0.23 98.9 ± 1.4 64.9 ± 1.4 .0472 ± 1.6 NVB207-8 0.30 225 120 0.55 93.7 ± 1.8 68.3 ± 1.9 .0503 ± 5.0 NVB207-9 0.05 1015 232 0.24 90.9 ± 1.4 70.6 ± 1.5 .0482 ± 2.4 NVB207-10 -0.01 1285 392 0.32 93.9 ± 1.4 68.3 ± 1.5 .0478 ± 2.0 NVB207-11 -0.41 357 94 0.27 92.4 ± 1.6 69.8 ± 1.7 .0446 ± 4.1 NVB207-12 -0.36 314 56 0.18 91.6 ± 1.6 70.3 ± 1.8 .0450 ± 4.4 NVB207-13 0.50 232 55 0.25 88.6 ± 1.7 72.1 ± 1.9 .0518 ± 4.8 NVB207-14 -0.03 419 127 0.31 90.5 ± 1.5 70.9 ± 1.7 .0476 ± 3.8

NVB-206: Granodiorite of Juniper Pass NVB206-1 0.66 152 76 0.52 94.8 ± 1.8 67.0 ± 1.9 .0532 ± 6.0 NVB206-2 0.50 346 183 0.55 91.8 ± 1.3 69.4 ± 1.4 .0518 ± 5.0 NVB206-3 0.07 224 66 0.31 94.2 ± 1.5 67.9 ± 1.6 .0485 ± 4.8 NVB206-4 0.89 1420 187 0.14 43.9 ± 0.6 145.0 ± 1.3 .0593 ± 4.9 NVB206-5 0.09 437 181 0.43 88.5 ± 1.2 72.2 ± 1.3 .0485 ± 3.5 NVB206-6 -0.44 230 112 0.51 90.9 ± 1.4 70.7 ± 1.5 .0444 ± 4.9 NVB206-7 -0.16 271 153 0.58 92.8 ± 1.4 69.1 ± 1.5 .0444 ± 4.4 NVB206-8 -0.02 475 186 0.41 91.7 ± 1.2 69.8 ± 1.3 .0477 ± 3.5 NVB206-9 0.34 364 147 0.42 89.1 ± 1.3 71.6 ± 1.4 .0505 ± 3.9 NVB206-10 0.43 253 76 0.31 92.8 ± 1.5 68.6 ± 1.6 .0513 ± 4.7 NVB206-11 0.65 1010 453 0.46 84.2 ± 1.0 75.5 ± 1.2 .0528 ± 2.6 NVB206-12 0.10 728 398 0.56 93.3 ± 1.1 68.6 ± 1.2 .0486 ± 2.8 NVB206-13 0.29 235 69 0.30 94.1 ± 1.5 67.8 ± 1.5 .0502 ± 4.7

NVB-286: Granodiorite in the Trinity Range 286-1 0.20 765 286 0.39 84.6 ± 0.7 75.6 ± 0.8 .0493 ± 2.9 286-2 0.64 557 376 0.70 89.9 ± 0.8 70.7 ± 0.9 .0529 ± 3.2 286-3 0.18 523 214 0.42 90.2 ± 0.8 70.8 ± 0.9 .0492 ± 3.3 286-4 0.13 954 497 0.54 90.7 ± 0.6 70.5 ± 0.7 .0488 ± 3.9 286-5 0.17 1140 200 0.18 90.2 ± 0.6 70.9 ± 0.6 .0492 ± 2.2 286-6 2.01 1040 278 0.28 92.1 ± 0.6 68.1 ± 0.6 .0638 ± 2.1 286-7 0.13 1833 681 0.38 90.6 ± 0.4 70.5 ± 0.5 .0489 ± 1.9 286-8 0.07 664 210 0.33 95.3 ± 0.9 67.1 ± 0.9 .0485 ± 3.2 286-9 2.39 408 116 0.29 88.3 ± 1.6 70.8 ± 1.5 .0667 ± 11.8 286-10 0.51 1999 164 0.08 138.9 ± 0.6 45.8 ± 0.5 .0528 ± 1.1 286-11 0.03 1576 1368 0.90 93.3 ± 0.5 68.5 ± 0.5 .0481 ± 1.9 286-12 0.03 574 258 0.46 92.2 ± 0.8 69.4 ± 0.9 .0481 ± 1.9 286-13 0.44 622 222 0.37 88.1 ± 0.8 72.3 ± 0.9 .0513 ± 3.2 286-14 0.17 942 624 0.68 90.1 ± 0.7 70.9 ± 0.8 .0491 ± 2.5 286-15 0.13 432 172 0.41 90.2 ± 1.1 70.9 ± 1.2 .0488 ± 3.8 133

TABLE A1. U-Pb SHRIMP ANALYTICAL DATA (CONTINUED)

Spot number Common 206Pb U Th 232Th/238U 206Pb/238U Age ± 1σ Total 238U/206Pb (± Total 207Pb/206Pb (± (%) (ppm) (ppm) (Ma) % err) % err)

NVB-212: Selenite Granodiorite NVB212-1 0.69 547 196 0.37 95.0 ± 1.5 67.1 ± 1.6 .0534 ± 2.9 NVB212-2 -0.37 622 154 0.26 94.2 ± 1.5 68.4 ± 1.6 .0450 ± 2.9 NVB212-3 0.15 1462 446 0.32 97.1 ± 1.4 66.0 ± 1.5 .0492 ± 1.8 NVB212-4 -0.07 479 110 0.24 95.6 ± 1.6 67.2 ± 1.6 .0474 ± 3.3 NVB212-5 0.16 3479 1485 0.44 98.6 ± 1.4 64.9 ± 1.4 .0493 ± 1.1 NVB212-6 4.01 531 118 0.23 96.3 ± 1.6 64.0 ± 1.6 .0798 ± 5.3 NVB212-9 0.37 615 185 0.31 98.8 ± 1.6 64.7 ± 1.6 .0509 ± 2.7 NVB212-10 0.06 686 153 0.23 97.2 ± 1.5 65.9 ± 1.5 .0484 ± 2.7 NVB212-11 0.11 1219 320 0.27 94.2 ± 1.4 68.0 ± 1.5 .0487 ± 2.1 NVB212-12 -0.26 784 205 0.27 99.7 ± 1.5 64.5 ± 1.5 .0459 ± 2.7 NVB212-13 -0.18 296 77 0.27 96.6 ± 1.7 66.5 ± 1.7 .0465 ± 4.2 NVB212-16 0.19 282 103 0.38 96.7 ± 1.7 66.2 ± 1.7 .0494 ± 4.0 NVB212-17 -0.03 1434 356 0.26 101.6 ± 1.5 63.1 ± 1.5 .0478 ± 1.8 NVB212-18 2.80 503 131 0.27 100.5 ± 1.6 62.1 ± 1.6 .0702 ± 2.6

NVB-208: Power Line Complex NVB208-1 -0.04 2244 236 0.11 110.0 ± 1.6 58.3 ± 1.4 .0479 ± 1.5 NVB208-2 -0.06 4127 1250 0.31 110.2 ± 1.6 58.2 ± 1.4 .0477 ± 1.1 NVB208-3 -0.05 908 857 0.98 105.8 ± 1.6 60.6 ± 1.5 .0478 ± 2.3 NVB208-4 4.26 624 157 0.26 102.7 ± 1.6 59.8 ± 1.5 .0819 ± 2.1 NVB208-5 0.31 753 169 0.23 99.8 ± 1.5 64.1 ± 1.5 .0504 ± 3.3 NVB208-6 -0.16 744 153 0.21 104.7 ± 1.7 61.4 ± 1.6 .0468 ± 2.9 NVB208-7 0.19 1218 311 0.26 105.2 ± 1.6 60.8 ± 1.5 .0496 ± 2.0 NVB208-8 0.11 766 179 0.24 105.5 ± 1.6 60.7 ± 1.5 .0490 ± 2.5 NVB208-9 0.08 1117 329 0.30 109.0 ± 1.6 58.8 ± 1.5 .0488 ± 2.1 NVB208-10 2.63 961 254 0.27 100.3 ± 1.5 62.2 ± 1.5 .0689 ± 1.8 NVB208-11 -0.10 4772 750 0.16 110.0 ± 1.5 58.3 ± 1.4 .0474 ± 1.0 NVB208-12 -0.05 555 121 0.23 106.6 ± 1.7 60.2 ± 1.6 .0477 ± 2.9 NVB208-13 0.06 690 169 0.25 103.9 ± 1.6 61.7 ± 1.5 .0486 ± 2.6 NVB208-14 0.23 666 177 0.27 104.0 ± 1.6 61.5 ± 1.5 .0499 ± 2.6

NVB280: Boudinaged pegmatite dike 280-1 0.06 14877 770 0.05 97.2 ± 0.6 65.8 ± 0.6 .0485 ± 1.0 280-3 0.15 1767 41 0.02 89.3 ± 0.5 71.6 ± 0.6 .0490 ± 2.1 280-5 0.66 8479 312 0.04 95.3 ± 0.2 66.7 ± 0.2 .0531 ± 0.9 280-6 2.56 6428 235 0.04 94.3 ± 0.4 66.2 ± 0.3 .0682 ± 3.6 134

CHAPTER 4: A MULTI-THERMOCHRONOMETER RECORD OF EROSIONAL EXHUMATION OF THE SIERRA NEVADA BATHOLITH IN THE NORTHWEST BASIN AND RANGE

Nicholas J. Van Buer, Trevor A. Dumitru, Martin J. Grove Department of Geological and Environmental Sciences, Stanford University 450 Serra Mall, Bldg. 320, Stanford, CA 94305-2115

Jeremy K. Hourigan Earth and Planetary Sciences Department, University of California, Santa Cruz Earth & Marine Sci., Santa Cruz, CA 95064

135

ABSTRACT

Understanding the Late Cretaceous to early Tertiary erosional history of the northern Sierra Nevada Batholith is important for understanding the paleogeography of the Sierra Nevada and that of the proposed “Nevadaplano” orogenic plateau that once extended east of the batholith. Basin and Range extension has dismembered the northernmost Sierra Nevada Batholith and overlying Tertiary strata, which were deposited unconformably on eroded rocks of the batholith. Large-offset normal faults bound tilted crustal sections whose structural relief provides exposure of the upper 7 km of the batholith. We take advantage of the ability to sample the batholith with depth to better determine its thermal history from the time of its intrusion in the Late Cretaceous to the development of the Tertiary unconformity using a combination of five medium- to low-temperature thermochronometers applied to nine paleodepth transects. This thermal history reflects post-emplacement cooling of the batholith followed by steadily- decreasing exhumation rates from the Late Cretaceous into the mid-Tertiary. The duration of erosional exhumation is correlated with the age of the batholithic rocks, lasting longer in the eastern part of the arc, suggesting that residual arc heat may have helped maintain high arc topography from the Late Cretaceous until mid-Tertiary time.

INTRODUCTION

The rugged Sierra Nevada, mostly underlain by Mesozoic batholithic rocks, is bound by normal faults on its eastern side that produce the impressive relative relief and steep slopes of its eastern escarpment (Fig. 4.1). The range gradually loses that relief northward and meets its topographic demise near Susanville, California, but rocks of the Nicholas J. Van Buer and Elizabeth L. Miller

136

126° W 124° W 122° W 120° W 118°118° W 116° W 114° W 112° W 110° W 42° N OR Marginal Idaho Sn Batholith ID 87 86 CA a terranes Sr/ Sr = 0.706 k ? e n ck 44° N Klamath Riv lai ra er P t t Range NV otspo 40° N Yellowstone h

Susanville area of 42° N Figs. 4.2, 4.3, 4.11

Sierra Nevada Batholith Reno 38° N LFTB B a s i n a n d UT San Francisco R a n g e 40° N

36° N Sevier Thrust Belt North

American 38° N AZ N craton 34° N Tehachapi

36° N Los Angeles Peninsular 32° N Ranges 500 km Batholith 34° N

122° W 120° W 118° W 116° W 114° W 112° W 110° W 108° W Figure 4.1. Rocks of the Cordilleran magmatic arc (white) form a nearly continuous belt across the western United States, though divided by Cenozoic events into the Sierra Nevada Batholith, the Peninsular Ranges Batholith, and the Idaho Batholith. The most continuous exposures occur in the Sierra Nevada mountains, between Tehachapi and Susanville. North of the initial 87 Sr/ 86 Sr line (dashed) the arc crosses from continental lithosphere into marginal terranes of more oceanic affinity. This transition roughly corresponds with the topographic demise of the Sierra Nevada Batholith as it enters the Basin and Range (surrounded by dotted lines), where batholithic outcrop is less continuous due to Cenozoic cover and extension. LFTB = Luning-Fencemaker Thrust Belt. Adapted from Van Buer and Miller (2010). 137 batholith continue across part of Basin and Range province in northwestern Nevada

(Figs. 4.1, 4.2; Van Buer et al., 2009; Van Buer and Miller, 2010). At present, there is substantial controversy about the fundamental mechanisms and exact timing of the modern uplift of the Sierra Nevada (see discussions in Wakabayashi and Sawyer, 2001;

Jones et al., 2004; Busby and Putirka, 2009). The physiographic nature of the ancestral

Sierra Nevada at the end of arc magmatism in the Late Cretaceous and its topographic development into the early Tertiary is less well-constrained and even more controversial.

Given the distinction between the modern Sierra Nevada mountain range versus the extent of the batholith itself, the degree to which uplift of the modern Sierra Nevada can be linked to its older intrusive history is debatable. On the other hand, arc magmatism almost certainly exerted a primary influence on Mesozoic topography, both during arc activity and immediately afterwards. This study exploits the 3-D views into the upper crust of the northernmost Sierra Nevada batholith afforded by Basin and Range normal faulting to decipher the preserved thermal history of the batholith from its Cretaceous intrusion to its uplift and erosion, which continued into the mid-Tertiary.

Many workers have suggested that the Sierra Nevada developed its most imposing topography in the Late Cretaceous (~ 85 Ma), during the end-stages of crustal shortening and emplacement of the batholith and that it has been slowly eroding since then, as documented best by thermochronologic studies (House et al., 1997, 1998, 2001;

Cecil et al., 2006), stable isotope studies (e.g., Mulch et al., 2006), and the development of a well-defined erosional surface in the northern part of the Sierra Nevada (Fig. 4.2).

Specifically, the ancestral Late Cretaceous-early Tertiary Sierra Nevada is thought to have formed the western margin of a broad, high plateau underlain by overthickened 138

Figure 4.2 (next page): Extent of the early- to mid-Tertiary unconformity in the northern Sierra Nevada Batholith and adjacent Basin and Range, showing subcrop geology. Squiggled lines, colored by basement type, represent outcrops of the unconformity, and the colored fields represent early Tertiary surface geology (distorted by later extension). LFTB = Luning-Fencemaker Thrust Belt; only the frontal thrusts of more extensive thrust systems are shown. Short, heavy black lines indicate the unconformity-controlled lines of section used in this study (Fig. 4.5). Modified from Van Buer et al. (2009) and Martin et al. (2010). 139

42° N CA NV (no data) AA’

100 km Mza BB’ EE’

Mzg King Lear CC’ Modoc Plateau Fm.

(no data) JTrb lPz lPz LFTB

GG’ DD’ Mza ?

FF’ ? Trs

uPz uPz HH’ LFTB uPz Mzg

JTrb

(no data) Golconda Thrust lPz

Mza

lPz II’ Trs N LFTB

uPz uPz Mzg Mza 38° lPz N 121° W Legend 117° W Ks Cretaceous sediments

Unconformity Mzg Granitic rocks of the Sierra Nevada Batholith (color coded JTrb Late Triassic to earliest Jurassic basinal terrane by basement) Trs Triassic shelf deposits Mza Major thrust front Mesozoic volcanic arc terranes (some Pz rocks locally) uPz Upper Paleozoic wall rocks and Golconda allochthon Figure 4.2 (continued) lPz Lower Paleozoic wall rocks and Roberts Mts. allochthon 140 crust built up behind the Sevier thrust belt to the east (Fig. 4.1; e.g. Allmendinger, 1992;

Chase et al., 1998; DeCelles, 2004; DeCelles and Coogan, 2006; Colgan and Henry,

2009). The exact extent and elevation of this plateau, termed the “Nevadaplano”

(DeCelles, 2004), as well as the duration of its existence, remain controversial. For example, Mulch et al. (2006) and Cassel et al. (2009) present stable isotope data that suggest the western slopes of the Sierra Nevada in the Eocene were as steep as those today, inconsistent with some of the accumulated evidence for young, Basin and Range related tilting, uplift and incision (e.g. Wakabayashi and Sawyer, 2001). Conversely, initial results from geophysical imaging of the delamination or dripping-off of dense lithosphere from the base of the batholith suggest that the loss of the dense lithospheric root of the batholith and its replacement by asthenospheric mantle did cause recent surface uplift, and was a necessary driver for the onset of Basin and Range faulting (e.g.,

Jones et al., 2004; Gilbert et al., 2007; Reeg, 2008).

Regarding the extent of the Nevadaplano, recent seismic experiments (Lerch et al., 2007, 2008a; Gashawbeza et al., 2008), coupled with studies of the magnitude and age of Basin and Range faulting (Colgan et al., 2004, 2006a, 2006b, 2008; Lerch et al.,

2008b, 2010; Fosdick and Colgan, 2008; Whitehill, 2009; Egger, 2010), conclude that the northern part of the batholith may never have been underlain by crust as thick or as low- velocity as that of the southern Sierra (Fig. 4.1). Evidence for a thin crust beneath the batholith contrasts with large inferred amounts of erosional stripping (about 3 to 7 km) and thus topographic uplift of the batholith prior to development of the unconformity

(Van Buer et al., 2009). During the later intrusive history of the batholith, marine Upper

Cretaceous sedimentary successions prograded across basement rocks of the Klamath 141

Mountains immediately to the west (Nilsen, 1984, 1993) and Eocene sequences exposed in the Warner Range, CA, eastern and northern Klamath Mountains and coastal Oregon, appear to be at least in part derived from the erosion of the batholith in Nevada (Figs. 4.1,

4.2, 4.3; McKnight, 1984; Wooden et al., 1999; Egger et al., 2009). Thus these disparate lines of evidence point to changes in overall elevations, crustal structure, and landform development across the northern part of the batholith in the Late Cretaceous or early

Cenozoic, suggesting a topographically subdued region (cf. Saleeby et al., 2009) with drainages carrying material westward and/or northward. This view is in distinct contrast to studies implying that an altiplano-style plateau extended to the northern Sierra Nevada batholith from the end of Cretaceous magmatism to the beginning of Basin and Range extension (e.g. Ducea, 2001; DeCelles, 2004; DeCelles and Coogan, 2006; Mulch et al.,

2006; Cassel et al., 2009).

To help constrain the late Cretaceous to early Tertiary paleogeography of the northernmost Sierra Nevada batholith, this study uses a combination of thermochronometers, including apatite fission track and U-Th/He, zircon U-Th/He,

39Ar/40Ar biotite and 39Ar/40Ar K-feldspar multi-diffusion domain modeling, to read the thermal signal of cooling related to the erosion and uplift to the surface of batholithic rocks.. The tilted crustal sections of the batholith in the Basin and Range offer an unparalled opportunity for thermochronologic studies that target the batholith’s post- intrusive cooling and erosional history up to the time of development of a widespread regional unconformity overlain by Eocene to Miocene strata preserved in many of the ranges as well as in the northern Sierra Nevada proper (Fig. 4.2; Van Buer et al., 2009).

This study is centered in NW Nevada (Fig. 4.1), but because its Mesozoic bedrock 142

42° OR N Bilk Creek R.

Pueblo Mts. A MODOC Pine PLATEAU Forest Warner Range (Colgan et al., 2006) (Colgan et al., R. B 2006; 2008) Black Rock Range E Santa Rosa R.

Jackson Range

Bloody Run Hills Winnemucca

NV CA

Seven East Range

GSelenite R. D Troughs Susanville R. F (Fosdick and Diamond Mountains Colgan, 2008) S I E R R A N E V A D A (Whitehill, 2009) Sahwave R. H

Nightingale R.

Reno

(Cecil et al., 2006)

Carson R. Pine 38° Nut Wassuk Range N (Surpless,Range 1999; Stockli et al, 2002) I

121° W 118° W Figure 4.3. Shaded relief map of the northern Sierra Nevada and northwestern Basin and Range, showing the locations of previous thermochronology samples (black dots, labeled with source). Lettered boxes show the areas sampled in this study (Fig. 4.4A - 4.4I); heavy-outlined boxes are also sites of previous work. 143 history is shared with that of the northern Sierra Nevada, the results of this study can be extrapolated to a much broader region. Although the focus of this study is on the earlier erosional history of the ancestral Sierra Nevada, the data also provide additional data with which to understand the more contemporary evolution of the batholith.

GEOLOGIC BACKGROUND

Cordilleran subduction-related arc magmatism was active in the Triassic and continued intermittently until the Late Cretaceous (and into the Paleocene north of the

Snake River Plain and southwards from Arizona), building a more or less continuous batholithic belt across the western edge of North America (Fig. 4.1). In the U.S.

Cordillera, the Cretaceous intrusive complexes that form part of this batholith (Figs. 4.1,

4.2) generally become younger and more felsic to the east across its ~100 km width (e.g.,

Evernden and Kistler, 1970, Hyndman, 1983; Silver et al., 1979), culminating in a particularly voluminous set of intrusions, often concentrically zoned, along its eastern edge at ca. 94-83 Ma (e.g., Kistler et al., 1986; Walawender et al., 1990; Coleman and

Glazner, 1997). Earlier Mesozoic intrusions are less consistently organized, partly due to tectonic disruption and younger intrusive history (e.g., Evernden and Kistler, 1970,

Saleeby and Busby-Spera, 1992). Cenozoic strike-slip and extensional faulting together with extensive Cenozoic volcanic cover has broken this magmatic arc into distinct map segments, including the Idaho Batholith, the Sierra Nevada Batholith, and the Peninsular

Ranges Batholith (Fig. 4.1; e.g. Barton et al., 1988). The metamorphic wall rocks of the northwestern Sierra Nevada and the Jurassic and early Cretaceous plutons that intrude them have been closely compared to rock units in the Klamath Mountains to the 144 northwest, but the eastern and younger Cretaceous part of the batholith diverges into the northwestern Basin and Range (Fig. 4.1, 4.2; Van Buer et al., 2009; Van Buer and Miller,

2010). Although disrupted by normal faulting and largely covered by Cenozoic deposits, batholithic rocks actually make up most of the basement outcrop in the NW corner of

Nevada, as revealed by the units exposed beneath the Tertiary unconformity (Fig. 4.2, e.g., Smith et al., 1971; Barton et al., 1988; Van Buer et al., 2009; Van Buer and Miller,

2010). These plutonic rocks include voluminous, mostly granodioritic intrusions associated with the ca. 90 Ma high-flux arc flare-up along the Sierran crest (Van Buer and Miller, 2010), which show strong petrological and geochemical similarities to coeval intrusions in the Sierra Nevada (Van Buer and Miller, 2010).

The Sierra Nevada Batholith exhibits pronounced changes from north to south, traditionally ascribed to the fundamental architecture of its country rocks: to the south, the batholith was emplaced across the truncated edge of the Cordilleran miogeocline, while to the north, it was emplaced into basinal, oceanic, and accreted island arc terranes

(Fig. 4.1, 4.2). The boundary between mostly continental and mostly oceanic rocks is approximated by the initial 87Sr/86Sr = 0.706 isopleth (Fig. 4.1; e.g., Gastil, 1975;

Saleeby, 1981; Kistler, 1990). In the study area, most of the wall rocks to the batholith are Upper Triassic to lowest Jurassic slaty to phyllitic basinal metasedimentary rocks, but

Paleozoic to early Mesozoic arc rocks occur in the Pine Forest and Wassuk Ranges (Fig.

4.2, 4.3; e.g., Wyld, 1996; Surpless, 1999; Van Buer et al., 2009). Triassic/Jurassic country rocks are locally metamorphosed to amphibolite-grade siliceous hornfels where intruded by batholithic rocks, and pelitic rocks locally contain andalusite and cordierite.

Where undisturbed by Cretaceous intrusives, the basinal rocks record Jurassic cleavage 145 development as well as southeast-vergent folding and thrusting related to the Luning-

Fencemaker thrust belt (Fig. 4.1, 4.2; e.g., Compton, 1960; Oldow, 1984; Wyld et al.,

2001). Cretaceous deformation is limited to relatively narrow (kilometer-scale) aureoles of plutons (e.g., Compton, 1960; Ciavarella and Wyld, 2008; Van Buer and Miller, 2010).

Lower Cretaceous terrestrial strata preserved in isolated half-grabens (Fig. 4.2) show no deformation at all—these rocks do not record deformation related to the Sevier orogeny to the east (Fig. 4.1; Martin et al., 2010; cf. DeCelles et al., 2004).

The change from continental to basinal terrane wall rocks also appears to coincide with the preservation of Cenozoic strata above an unconformity developed across the

Mesozoic arc: in the northern Sierra Nevada and the Basin and Range Province, the batholith displays a well-developed erosional surface overlain by Tertiary strata as old as

Eocene and as young as about 8 Ma (Stewart, 1998; Wakabayashi and Sawyer, 2001;

Henry, 2008; Van Buer et al., 2009) whereas to the south, evidence for the unconformity and its overlying rocks is mostly absent (Fig. 4.1, 4.2). This unconformity is referred to as the “Tertiary unconformity” in this paper for brevity, although it actually represents a period of erosion that began during final intrusion of the batholith. Tertiary strata in the study area are primarily volcanic rocks, although sedimentary deposits are present locally. The oldest but least volumetrically significant rocks above the unconformity are

Eocene in age. In the Sierra Nevada, the basal unit, which occurs only in paleochannels, is the auriferous gravel of California gold rush fame (e.g., Lindgren, 1911; Yeend, 1974;

Cassel and Graham, 2011), constrained as upper Lower Eocene by plant fossils (e.g.,

MacGintie, 1941; Wolfe, 1994). Additionally, 38-35 Ma Eocene basalts cap low-relief basement topography in the very northwestern part of the study area, for example in the 146

Pine Forest Range (Fig. 4.4A; Colgan et al., 2006a) and the Black Rock Range (Fig.

4.4E; Lerch et al., 2008b). These ranges also expose Oligocene age lavas; elsewhere in the study area Oligocene units are predominantly 32-24 Ma ash-flow tuffs confined to paleochannels which are 5-10 km wide and up to 1.6 km deep (Henry, 2008). The presence of a smattering of mid-Tertiary units spanning a broad range of ages seems to indicate a long period of relatively minimal net erosion or deposition between the Eocene and the Miocene (cf. Van Buer et al, 2009).

Miocene strata, including ignimbrite deposits, basalt flows, and locally more silicic flows, often interbedded with fluvial, alluvial and lacustrine strata, are more widespread and are a kilometer or so thick where preserved (Table 4.1; e.g., Noble et al.,

1970; Colgan et al., 2006b). The Lower Miocene strata in the northwest part of the study area include 17-15 Ma flood basalts and peralkaline rhyolites traditionally associated with impingement of the Yellowstone hotspot (Fig 4.1), and predate extensional faulting

(e.g., Rytuba and McKee, 1984; Pierce and Morgan, 1992; Colgan et al., 2004; Coble et al., 2007). Some of the younger strata, however, locally are reported to exhibit evidence for growth faulting (e.g., Whitehill, 2009). The oldest angular unconformity observed is in the Wassuk Range, between the Lincoln Flat Andesite and the overlying Wassuk group, dated at 15 and 14.5 Ma, respectively (Fig. 4.5I; Surpless et al., 2002).

Thermochronological constraints suggest major normal faulting began at ~14.5 Ma in the southern half of the study area (e.g., Wassuk Range and environs; Surpless et al., 2002;

Sahwave Range; Whitehill, 2009), but in the northern half of the area, faulting began 12-

11 m.y. ago (Fig. 4.3; Colgan et. al. 2004, 2006ab, 2008). Extension has continued at lesser slip rates and perhaps episodically, to the present. Holocene scarps along range- 147

Symbols contact A. PINE FOREST RANGE 41° 50’ N normal faults: A inferred concealed sample locality (this study) sample locality A’ (previous study)

41° 45’ N 118° 45’ W 118° 40’ W 118° 35’ W (modified from Colgan, 2005; Colgan et al., 2010) samples B. SANTA ROSA RANGE not in Fig. 4.7

Sawtooth Stock 41° 35’ N B Santa Rosa pluton (105 Ma)

B’

41° 30’ N 117° 45’ W 117° 40’ W 117° 35’ W (modified from Compton, 1960) D. SEVEN TROUGHS RANGE

C. BLOODY RUN HILLS 40° 35’ N C N D Bloody Run Stock D’

41° 10’ N

C’ 5 km 117° 45’ W 118° 50’ W 118° 45’ W (modified from Compton, 1960) (this study; Jennings, 1977) Figure 4.4 (continued on next page): Simplified geologic maps of the nine ranges sampled for thermochronology, showing the lines of the cross sections in Figure 4.5 as well as previous thermochronology samples (black dots) and our samples (open circles). All maps at same scale (1:180,000). Explanation of map units is on second page. 148

41° 30’ N E. BLACK ROCK RANGE F. DIAMOND MOUNTAINS F’ E H o n e y L a k e F a u l t Z o n e 40° 20’ N

E’

41° 25’ N 118° 55’ W 118° 50’ W (modified from Lerch et al., 2008) Thompson Pk.

F 40° 15’ N 120° 35’ W 120° 30’ W (modified from Grose and Porro, 1989) G. SELENITE RANGE H. SAHWAVE RANGE 40° 05’ N

H H’

40° 35’ N G’ (this study) 119° 15’ W 119° 10’ W 119° 05’ W 119° W (modified from Van Buer and Miller, 2010) I. WASSUK RANGE Geologic Units I’ 38° 50’ N Quaternary cover

Miocene (16-9 Ma) mostly tuff

Miocene (17-15 Ma) mostly basalt

Oligocene and I Eocene volcanics 38° 45’ N 118° 55’ W 118° 50’ W 118° 45’ W (modified from Surpless, 1999) Figure 4.4 (continued) Tertiary intrusions N Cretaceous intrusions 5 km Pre-Cretaceous metamorphic rocks 149

TABLE 4.1. GEOLOGICAL COMPARISON OF THE SAMPLE TRANSECTS

Location Basement age Oldest Tertiary Depth of Tertiary strata (Ma, 2σ) strata (Ma) section* (m) thickness† (m)

A. Pine Forest Range ~108 30 4300 1100 B. Santa Rosa Range 105.1 ± 1.1 17 3800 800 C. Bloody Run Hills 97.7 ± 0.7 17? 1800 800 D. Seven Troughs Range 102.9 ± 1.5 30? 1000 >500 E. Black Rock Range 112.3 ± 1.4 35 350 1000 F. Diamond Mountains ~105 16 (locally) 800 >300 G. Selenite Range 96.4 ± 1.2 18? 700 >200 H. Sahwave Range ~93 18 2800 300 - 800 I. Wassuk Range ~85? 27 7200 800

*Maximum well-constrained paleodepth beneath basal Tertiary unconformity †Maximum exposed thickness of Tertiary strata (affects late Tertiary reheating) All data sources listed in the text. 150

Symbols projected PINE FOREST RANGE unconformity

3A A’3 1 km 4 km depth below F F 2 km F,Z break in section Z,K unconformity 2 3 km 2

1 ? 4 km 1 km sample locality km (this study) s. l. ? s. l. sample locality -1 -1 Left: from mapping by Colgan et al., 2010; Right: modified from Colgan et al. 2006b (previous study) -2 -2 ? samples ? SANTA ROSA RANGE ? not in Fig. 4.7 ? unconformity dip might be steeper than projected. B A,F,Z B’3 3 F F Z 1 km 2 2 K 2 km 4 km 3 km 1 1 km

km s. l. s. l. break in section

-1 -1 Left: modified from Colgan et al., 2006b; Right: from mapping by Compton, 1960 -2 -2

BLOODY RUN HILLS SEVEN TROUGHS RANGE C Bloody Run Peak C’ D D’ 3 F 3 3 A,F,Z,K,B 3 A,F A,F F F,Z,K,B F F A,F 2 2 2 2 1 km 1 km

2 km km 1 1 1 ? 1

km km ? s. l. s. l. s. l. s. l.

-1 -1 -1 -1 From mapping by Compton, 1960 Previous thermochronology samples from Colgan et al., 2006b This study -2 -2 -2 -2

BLACK ROCK RANGE 5 km E E’ 3 3

F 2 F,B 2

1 1 km 1 km

km s. l. s. l. ? -1 -1 From mapping by Lerch et al., 2008 -2 -2 Figure 4.5 (continued on next page): Cross sections of ranges discussed in text showing sample paleodepths beneath the Tertiary unconformity. Data in Fig. 4.7 (not grey) labeled by type: A - apatite U-Th/He, Z - zircon U-Th/He, F - apatite fission track, K - K-feldspar 40 Ar/ 39 Ar, B - biotite 40 Ar/ 39 Ar. Same units and symbols as Fig. 4.4. No vertical exaggeration. Data sources shown on individual cross sections. 151

DIAMOND MOUNTAINS Geologic Units F F’ 3 3 A,F Quaternary cover F 2 A,F 2 1 km Miocene (16-9 Ma)

km 1 ? 1

km mostly tuff s. l. s. l.

-1 -1 Miocene (17-15 Ma) From mapping by Grose and Porro, 1989 mostly basalt -2 -2 5 km Oligocene and SELENITE RANGE Eocene volcanics G’ 3G 3 A,F F 2 2 Tertiary intrusions

1 1 km 1 km

km Cretaceous intrusions s. l. s. l. bend -1 in -1 Pre-Cretaceous This study section metamorphic rocks -2 -2

SAHWAVE RANGE H 3 H’3 F,A F 1 km 2 F,Z F F,Z,K 2

1 2 km 3 km 1 km km ? s. l. s. l.

-1 bend -1 in section Modified from Van Buer, submitted -2 -2

WASSUK RANGE ? I Z I’ 3 1 km 2 km 3 F 3 km 4 kmF 5 km 2 6 km K 2

km 1 ? ? 7 km 1

km s. l. s. l.

-1 -1 From mapping by Surpless, 1999 -2 -2 Figure 4.5 (continued) 152 bounding faults are common, although not ubiquitous (e.g., Surpless, 1999; Wesnousky et al., 2005). Normal faulting has created ranges that are tilted structural blocks, which typically expose basement rocks along the bounding normal fault, while tilted Tertiary strata crop out on the other side of the tilt-block (Fig 4.4, 4.5). Where basement structures have been carefully mapped, there is no evidence of faults younger than mid-

Cretaceous and older than the basal Tertiary strata (Fig. 4.4, 4.5; e.g., Compton, 1960;

Wyld, 1996; Van Buer, submitted).

Exposure of the Late Cretaceous batholith at the surface before deposition of early- to mid-Tertiary strata (Fig. 4.2) implies a substantial amount of erosion between the Late Cretaceous and the early Tertiary. The general absence of Cretaceous volcanic rocks or caldera structures suggests that there has been > 3 km of erosion since arc magmatism ceased (e.g., Van Buer et al., 2009). The presence of the low-pressure aluminosilicate polymorph andalusite in contact aureoles of Cretaceous plutons 4 km structurally below the unconformity suggests there has been no more than about 7 km of erosion (Compton, 1960; Colgan et al., 2006b). We use thermochronologic data to obtain a more detailed history of this erosional exhumation.

USING THERMOCHRONOLOGY TO CONSTRAIN EXHUMATION

Where crustal cooling is accomplished by surface erosion, low- temperature thermochronometers can be used to determine long-term erosional exhumation rates (e.g., Wagner and Reimer, 1972; House et al., 1997; Cecil et al., 2006.)

Because different thermochronometers close at different temperatures, they are able to record passage of rocks through different depths in the earth (and/or changes in the 153 geothermal gradient). The temperature/depth range in which a thermochronometer begins recording is known as its partial retention zone (PRZ; or partial annealing zone, PAZ, in the case of fission-track analysis). Using multiple chronometers in rocks sampled from an intact crustal section can provide a fairly continuous record of thermal history, from intrusive cooling to erosional and tectonic exhumation (e.g., Armstrong et al., 2003;

Grove et al., 2003; Saleeby et al., 2008). For example, within upper crustal rocks subject to a geothermal gradient of 25–30 °C, 40Ar/39Ar dating of biotite (nominal closure temperature ~350 °C) will likely record intrusive cooling, whereas erosional and tectonic exhumation are recorded at lower temperatures: K-feldspar 40Ar/39Ar diffusion domain analysis can record passage through depths of ca. 6-14 km (~150-350 °C), zircon U-

Th/He through ca. 5-9 km (~160-230 °C), apatite fission track through ca. 2-5 km (~60-

120 °C), and apatite U-Th/He through ca. 1-3 km (~45-85 °C). These techniques are discussed in general below, but the detailed analytical procedures used in this study are documented in Appendices B, C, and D, for apatite fission track (AFT), apatite and zircon U-Th/He, and biotite and K-feldspar 39Ar/40Ar, respectively.

All thermochronometers fundamentally rely on the competition between radioactive decay and thermally-activated diffusion/reaction kinetics that can be modeled using the Arrhenius law: k = Ae-Q/RT, where k is the rate of daughter-product loss, A is a constant, Q is the activation energy (for diffusion or annealing), R is the gas constant, and

T is the temperature. These “clock-resetting” kinetics vary between different thermochronometers, such that each records the time of cooling through a certain temperature interval (which depends on cooling rate). Noble gas thermochronometers rely on thermally activated volume diffusion loss of argon or helium produced by the 154 radioactive decay of potassium or uranium/thorium, respectively. For example, as an apatite crystal passes upwards through the base of the He partial retention zone, cooling below ~85°C, it will begin to accumulate radiogenic He faster than He is lost by diffusion. Once this crystal rises above the top of the He PRZ, cooling below ~45°C, apatite will retain effectively all of its radiogenic He (except that lost via α-ejection at the moment of decay). As a result, the apatite will record an apparent age somewhere between the times it entered and left the PRZ (depending in detail on the cooling history and other factors such as size and radiation damage; e.g. Farley et al., 1996; Flowers et al., 2009; see Appendix C). Fission track dating relies on the balance between production of damage tracks in a crystal lattice by the spontaneous fission of 238U and annealing of the crystal lattice. Fission tracks have been shown to anneal inwards from the ends of the tracks, such that they grow shorter over time (at a rate that increases with temperature); therefore, careful measurement and modeling of fission track lengths can be used to estimate cooling rates and evaluate possible cooling paths (e.g., Gleadow et al., 1986;

Green et al., 1989; Ketcham, 2005; Appendix B). Similarly, more advanced analysis is possible with K-feldspar 39Ar/40Ar dating, where careful step-heating can be used to determine the argon-loss Arrhenius parameters for each sample, under the assumption that it contains a range of diffusion domain sizes sensitive to different temperatures

(Lovera et al., 1997, 2002; Appendix D). The physical origin of this multi-diffusion- domain (MDD) behavior in K-feldspar remains obscure in detail, but continuous thermal histories modeled from measured diffusion parameters have been empirically shown to be robust (Lovera et al., 1997, 2002).

Thermochronometry data are most easily modeled and interpreted when 155 monotonic cooling through the PRZ can be inferred. For example, if cooling proceeds roughly linearly with inverse temperature, the age can be related to a particular closure temperature after the method of Dodson (1973). Fortunately, the erosional exhumation studied here can be reasonably assumed to have caused monotonic cooling from the time of emplacement of the batholith until the unconformable deposition of Tertiary strata above the batholith. The deposition of ~1 km of Tertiary cover, mostly 18-14 Ma (Table

4.1), would have temporarily reversed this monotonic cooling trend, resetting some thermochronometers that had previously risen past their partial retention/annealing zone

(e.g., Colgan et al., 2006b). This part of the thermal history is, however, at least roughly constrained by the ages and thicknesses of the overlying Tertiary strata preserved today

(Table 4.1). Thermochronometers can be reset by Tertiary intrusions as well, but this problem can largely be avoided by sampling in areas distant from mapped Tertiary intrusions, which are generally small or rare near the transects that form part of this study

(Fig. 4.4), and thus are expected to have spatially limited thermal effects at the paleodepths represented by our sample transects (Table 4.1). Very near the unconformity, thermochronometers may be reset by overlying Tertiary ignimbrites and lava flows, but only very thick cooling units would be expected to cause substantial heating more than a few tens of meters below the surface. Ages of thermochronometers that reside in their partial retention/annealing zone for extended periods of time can be difficult to interpret without additional sources of information. This study addresses the problem by using several different thermochronometers together, by using the additional cooling information available from apatite fission-track length modeling and K-feldspar

MDD modeling, and by comparing data from the full range of paleodepths exposed in 156 each transect.

Although apatite and zircon U-Th/He and fission track data from the northern

Sierra Nevada and the NW Basin and Range exist (Fig. 4.3; e. g., Surpless, 1999; Stockli et al., 2002; Cecil et al., 2006; Colgan et al., 2006ab; Whitehill, 2008), these data are far from adequate in terms of documenting the detailed cooling history of batholithic rocks in this region. With the exception of Cecil et al. (2006), these prior studies were optimized to establish the age of onset of Basin and Range faulting, not the earlier cooling and exhumation history of the batholith, and there are > 200 km gaps between these prior studies (Fig. 4.3). Our sampling strategy takes full advantage of the existing data, and concentrates on areas less heavily sampled by earlier work, such as just beneath the unconformity, where the history of erosional exhumation can be best documented

(Fig. 4.4, 4.5). Existing data from near the unconformity suggest significant erosional exhumation was taking place until at least ca. 90–60 Ma, continuing later in the east than in the west (Cecil et al., 2006; Colgan et al., 2006b).

In addition to new thermochronology work, three new SHRIMP U-Pb ages were measured where the intrusive age of sampled rocks was not precisely known (Table 4.1;

Figure 4.6); see Appendix E for the analytical procedure and complete data set.

Sampling Strategy

In order to place more thorough constraints on the regional history of exhumation of the batholith, we applied the five thermochronometers discussed above to nine transects beneath the Tertiary unconformity (Fig. 4.3, 4.4, 4.5). Our strategy was to collect samples for thermochronology from the broadest range of depths available that 157

238 206 207 206 238 U/ Pb Pb corrected Pb/ U ages 60 65 70 75

NVB-287 Seven Troughs Range .08 110 .07 105 .06

.05 100 .04 102.9 ± 1.5 Ma

115 DD-8 Diamond Mountains .06

.05 Pb

110 206

.04 Age (Ma) Age

105 Pb/ .03 207 .02 100 104.5 ± 1.3 Ma SL-11 Selenite Range .06 100 .05

95 .04

.03 207Pb/206Pb 90 96.4 ± 1.2 Ma 207Pb/206Pb .02 2σ errors shown in both types of plot. 110 100 90 Age on concordia (Ma)

Figure 4.6. SHRIMP U-Pb results. At left, selected spot ages used for weighted-mean ages are shown by solid bars. Rejected spot ages are shown by empty bars. Weighted averages are shown by gray lines. Diagrams on right are inverse concordia plots, showing accepted spot analyses with solid symbols and rejected spot analyses with empty symbols. Ages along concordia (gray line with tick marks) are shown on the corresponding tick marks at bottom. All errors are 2σ. 158 preserve a record of pre-unconformity exhumation. The broad range of paleodepths necessary to achieve these goals is provided by the tilted crustal sections represented in normal fault blocks of the Basin and Range and is constrained by our ability to directly reference the depth of samples and their cooling histories to the overlying Tertiary unconformity (Fig. 4.5; e.g., Colgan et al., 2004, 2006ab).

Several samples were collected in each of nine structural sections in the northwestern Basin and Range, located about 70 km apart (Fig. 4.3). Good structural control is already available at several carefully mapped locations, such as where detailed thermochronologic studies have been previously carried out (Colgan et al., 2004, 2006ab;

Stockli et al., 2002, Whitehill, 2008), but new geologic mapping was undertaken in the

Seven Troughs and Selenite Ranges to constrain the depth of samples beneath the unconformity (Fig. 4.3, 4.4D, 4.4G). Every transect consisted of structural sections through the Cretaceous batholith (Fig. 4.2), both to ensure a good supply of the minerals required for our analysis (apatite, zircon, biotite, and K-feldspar) and also to avoid complications related to multiple intrusion of the pre-batholithic rocks. New cross- sections have been drawn along the sample transects chosen for this study, demonstrating the paleodepths of our samples beneath the Tertiary erosion surface represented by the basal Tertiary unconformity (Fig 4.5). Because this study is focused on the history of erosion before deposition of Tertiary strata, we emphasize the paleodepth beneath the basal Tertiary unconformity (Du in Table 4.2). The thickness of the overlying Tertiary section, however, contributes to a correspondingly greater paleodepth by the mid-

Miocene, prior to extensional faulting. This pre-extensional paleodepth (Ds of Colgan et al., 2006b) is important for understanding which samples may have been partially reset 159

TABLE 4.2. SAMPLE LOCALITY AND AGE DATA

Sample data Ages * k b Sample Latitude Longitude Du Apatite He Apatite FT mean FT length Zircon He K-spar /Biotite Ar number (°N) (°W) (m) (Ma ± 1σ) (Ma ± 1σ) (µm ± 1σ) (Ma ± 1σ) (Ma ± 2σ)

A. Pine Forest Range PF-01 41° 48’ 53” 118° 41’ 40” 90 70.9 ± 3.4 13.17 ± 0.09 75 ± 15 JC01-PF14 41° 45’ 49” 118° 40’ 23” 230 75.0 ± 2.4† 13.52 ± 0.09† JC02-PF44 41° 34’ 00” 118° 45’ 20” 400 79.3 ± 4.0† 12.93 ± 0.09† JC01-PF13 41° 45’ 51” 118° 39’ 48” 520 49.4 ± 3.9† 83.1 ± 2.8† 13.61 ± 0.09† JC01-PF12 41° 45’ 53” 118° 39’ 14” 940 19.8 ± 0.6† 39.2 ± 1.9† 13.43 ± 0.10† JC01-PF11 41° 45’ 51” 118° 38’ 40” 1320 12.4 ± 0.5† 21.8 ± 1.1† 14.67 ± 0.12† JC01-PF10 41° 45’ 51” 118° 38’ 11” 1710 9.6 ± 0.7† 33.1 ± 1.7† 12.67 ± 0.11† JC01-PF9 41° 45’ 49” 118° 37’ 46” 2010 10.2 ± 0.5† 33.7 ± 1.8† 12.57 ± 0.14† JC01-PF8 41° 45’ 55” 118° 37’ 17” 2470 8.6 ± 0.5† 24.7 ± 1.9† 12.10 ± 0.31† JC01-PF7 41° 45’ 55” 118° 36’ 58” 2740 7.6 ± 0.3† 14.0 ± 1.2† 13.08 ± 0.22† JC01-PF6 41° 45’ 49” 118° 36’ 41” 2960 7.0 ± 0.6† 12.0 ± 1.0† 13.76 ± 0.38† JC01-PF5 41° 45’ 47” 118° 36’ 21” 3240 7.0 ± 0.9† 11.1 ± 1.0† 14.03 ± 0.21† JC01-PF4 41° 45’ 45” 118° 36’ 06” 3440 6.7 ± 0.3† 10.8 ± 1.0† 14.05 ± 0.51† JC01-PF3 41° 45’ 42” 118° 35’ 54” 3630 5.8 ± 0.9† 9.3 ± 1.0† 13.83 ± 0.37† JC01-PF2 41° 45’ 36” 118° 35’ 41” 3780 5.8 ± 0.9† 11.0 ± 1.0† 13.31 ± 0.28† JC00-PF1/PF-11 41° 46’ 10” 118° 35’ 27” 3810 5.2 ± 0.6† 8.1 ± 1.1† 13.85 ± 0.45† 31.0 ± 8.7 84.5 ± 2.0k JC02-PF47 41° 47’ 47” 118° 35’ 28” 4110 4.6 ± 1.2† 7.3 ± 1.0† JC02-PF48 41° 43’ 47” 118° 35’ 12” 4270 3.3 ± 0.5† 7.0 ± 0.7† 13.85 ± 0.45†

B. Santa Rosa Range JC00-SR4 41° 34’ 50” 117° 37’ 30” 150 50.8 ± 1.5† 77.1 ± 3.3 13.89 ± 0.12† SR-20 41° 32’ 28” 117° 39’ 23” 170 51.6 ± 6.3 68.9 ± 3.9 13.67 ± 0.12 55 ± 15 SR-22 41° 32’ 40” 117° 39’ 50” 560 49.1 ± 6.8 59.7 ± 3.7 13.42 ± 0.09 JC00-SR9 41° 34’ 58” 117° 39’ 13” 790 37.4 ± 6.3† 70.3 ± 4.5 13.30 ± 0.13† SR-24 41° 32’ 46” 117° 40’ 12” 980 60.9 ± 3.4 13.18 ± 0.17 JC00-SR13 41° 34’ 48” 117° 40’ 48” 1040 21.3 ± 7.5† 54.0 ± 3.0 13.82 ± 0.14† SR-26 41° 32’ 55” 117° 40’ 41” 1390 49.9 ± 2.8 12.77 ± 0.33 JC00-SR15 41° 34’ 50” 117° 41’ 44” 1570 19.4 ± 3.0† 52.8 ± 2.5 13.38 ± 0.12† JC02-SR49 41° 26’ 08” 117° 45’ 31” 2300 27 ± 20† 15.3 ± 1.6 13.10 ± 0.29† SR-15 41° 32’ 34” 117° 42’ 57” 2520 17.5 ± 2.7 12.96 ± 0.20 55.6 ± 7.7 JC00-SR17 41° 34’ 48” 117° 43’ 53” 2560 9.6 ± 0.8 14.28 ± 0.41† JC00-SR50 41° 34’ 44” 117° 44’ 01” 2620 12.5 ± 2.1† 11.4 ± 0.9 13.85 ± 0.59† JC00-SR51 41° 34’ 44” 117° 44’ 17” 2740 10.5 ± 0.8 13.25 ± 0.34† JC00-SR52 41° 34’ 56” 117° 44’ 30” 3030 13.2 ± 1.4 13.77 ± 0.31† JC00-SR19 41° 35’ 01” 117° 44’ 55” 3480 8.1 ± 0.9 12.00 ± 1.25† JC00-SR22 41° 34’ 55” 117° 45’ 08” 3630 8.8 ± 0.3† 8.2 ± 0.6 14.01 ± 0.12† JC00-SR29/SR-17 41° 34’ 08” 117° 45’ 34” 3930 13 ± 3† 6.4 ± 0.8 13.88 ± 0.25† 48.2 ± 0.4k

C. Bloody Run Hills BN-01 41° 09’ 51” 117° 43’ 58” 160 40.1 ± 1.7 67.4 ± 2.9 13.13 ± 0.13 59.0 ± 3.4 BN-03 41° 10’ 24” 117° 46’ 01” 100 58.5 ± 3.0 11.70 ± 0.23 JC02-BR2 41° 10’ 07” 117° 44’ 30” 300 33.6 ± 4.5† 55.0 ± 2.8 13.35 ± 0.12† BN-05 41° 10’ 27” 117° 46’ 32” 590 41.8 ± 6.6 51.6 ± 2.4 12.26 ± 0.19 BN-07 41° 10’ 44” 117° 46’ 39” 1000 49.4 ± 1.8 11.22 ± 0.19 BN-09 41° 11’ 05” 117° 47’ 19” 1370 47.9 ± 2.1 12.16 ± 0.15 JC02-BR3 41° 11’ 03” 117° 47’ 46” 1660 36.3 ± 2.6† 43.0 ± 2.8 11.46 ± 0.20† BN-11 41° 11’ 20” 117° 47’ 48” 1730 46.6 ± 3.6 12.09 ± 0.19 59.1 ± 1.5 85.8±0.4k/94.4±1.9b

* Du is paleodepth beneath the basal Tertiary unconformity †Data from Colgan et al., 2006 160

TABLE 4.2. SAMPLE LOCALITY AND AGE DATA (Continued)

Sample data Ages * Sample Latitude Longitude Du Apatite He Apatite FT mean FT length Zircon He K-spar/Biotite Ar number (°N) (°W) (m) (Ma ± 1σ) (Ma ± 1σ) (µm ± 1σ) (Ma ± 1σ) (Ma ± 2σ)

D. Seven Troughs Range 7T-1A 40° 33’ 32” 118° 47’ 02” 30 16.5 ± 1.1 14.09 ± 0.74 7T-3 40° 33’ 52” 118° 48’ 07” 450 43.0 ± 1.5 64.5 ± 2.4 12.57 ± 0.20 7T-5 40° 33’ 45” 118° 49’ 22” 790 42.2 ± 2.7 67.6 ± 2.1 12.85 ± 0.20 70.2 ± 3.9 86.2±1.0k/89.3±1.5b NVB-287 40° 32’ 27” 118° 48’ 35” 790

E. Black Rock Range NBR-3 41° 27’ 59” 118° 52’ 34” 30 77.2 ± 2.1 13.77 ± 0.08 JC02-BL5 41° 28’ 06” 118° 52’ 32” 90 56 ± 12† 87.0 ± 3.2 13.89 ± 0.07† JC02-BL6 41° 27’ 54” 118° 52’ 16” 190 48.9 ± 1.5† 80.2 ± 3.7 11.85 ± 0.07† JC02-BL7 41° 27’ 46” 118° 52’ 06” 220 43 ± 24† 78.7 ± 3.2 14.06 ± 0.07† NBR-1 41° 27’ 32” 118° 52’ 15” 320 79.1 ± 2.3 13.61 ± 0.10 105.3 ± 1.3b

F. Diamond Mountains DD-1 40° 17’ 26” 120° 35’ 13” 30 50 ± 17 89.3 ± 5.2 DD-5 40° 17’ 53” 120° 35’ 18” 560 84.3 ± 4.5 13.75 ± 0.11 DD-8 40° 18’ 27” 120° 35’ 15” 780 52.5 ± 6.3 74.8 ± 2.6 13.40 ± 0.11

G. Selenite Range SL-11 40° 36’ 06” 119° 15’ 25” 80 50 ± 11 62.3 ± 3.0 13.36 ± 0.12 SL-13 40° 35’ 43” 119° 15’ 18” 400 59.8 ± 2.9 13.43 ± 0.10

H. Sahwave Range SE-01 40° 02’ 52” 119° 08’ 07” 60 66.3 ± 3.0 12.14 ± 0.29 67.3 ± 8.3 SE-04 40° 02’ 51” 119° 07’ 01” 350 33 ± 14 61.2 ± 2.7 12.42 ± 0.12 SE-06 40° 02’ 42” 119° 05’ 42” 900 55.1 ± 1.9 12.41 ± 0.16 SE-08 40° 02’ 24” 119° 04’ 54” 1560 53.5 ± 2.1 12.84 ± 0.12 SE-10 40° 02’ 25” 119° 04’ 02” 2180 35.8 ± 1.6 12.51 ± 0.23 SE-12 40° 02’ 32” 119° 03’ 24” 2760 23.6 ± 1.3 12.73 ± 0.15 56 ± 14 81.0 ± 0.4k

I. Wassuk Range WK-01 38° 47’ 27” 118° 52’ 05” 70 124.5 ± 20.1 96BS-11.5c 38° 47’ 00” 118° 51’ 09” 540 50.8 ± 4.7§ 12.26 ± 0.29§ 96BS-11.5a 38° 47’ 08” 118° 50’ 20” 810 44.6 ± 3.5§ 12.12 ± 0.17§ 96BS-11.5 38° 47’ 05” 118° 49’ 59” 1710 31.2 ± 2.3§ 50.1 ± 5.0§ 12.15 ± 0.20§ 96BS-11.4c 38° 47’ 04” 118° 49’ 43” 2060 42.7 ± 5.2§ 12.40 ± 0.36§ 96BS-11.4b 38° 47’ 04” 118° 49’ 22” 2520 15.5 ± 2.0§ 13.75 ± 0.14§ WK-13 38° 48’ 14” 118° 50’ 17” 2870 49.5 ± 3.8 96BS-11.4a 38° 47’ 03” 118° 48’ 57” 3040 13.9 ± 2.4§ 13.74 ± 0.21§ 96BS-11.4 38° 47’ 01” 118° 48’ 25” 3800 14.5 ± 1.6§ 15.1 ± 1.8§ 13.75 ± 0.19§ 96BS-11.3a 38° 46’ 52” 118° 48’ 02” 4640 14.9 ± 1.0§ 13.91 ± 0.11§ 96BS-11.3 38° 46’ 52” 118° 47’ 36” 5470 9.9 ± 0.9§ 16.3 ± 1.4§ 13.76 ± 0.35§ 96BS-11.2b 38° 46’ 43” 118° 47’ 11” 5610 14.8 ± 1.2§ 13.48 ± 0.14§ 96BS-11.2a 38° 46’ 29” 118° 47’ 07” 5740 14.5 ± 1.2§ 13.15 ± 0.14§ 96BS-11.2 38° 46’ 18” 118° 46’ 56” 5920 7.0 ± 1.0§ 14.3 ± 1.2§ 13.62 ± 0.12§ 96BS-11.1b 38° 46’ 09” 118° 46’ 35” 6150 14.7 ± 1.2§ 13.48 ± 0.30§ 96BS-11.1a 38° 45’ 59” 118° 46’ 21” 6360 6.7 ± 0.6§ 12.8 ± 1.2§ 13.46 ± 0.14§ 96BS-11.1 38° 45’ 53” 118° 45’ 52” 7100 3.9 ± 0.8§ 10.1 ± 1.4§ 13.09 ± 0.15§ WK-15 38° 46’ 48” 118° 46’ 48” 7200 43.1 ± 0.3k

* Du is paleodepth beneath the basal Tertiary unconformity †Data from Colgan et al., 2006 §Data from Stockli et al., 2002 161 by the rise in temperature—and the corresponding rise of the partial retention/annealing zone—during the deposition of Tertiary strata.

THERMOCHRONOLOGY RESULTS

Thermochronology results for each of the nine transects are summarized in Table

4.2 and Figure 4.7. Among the thermochronometers used, apatite fission track age and length data (Tables 4.2, B1) provide the most useful and consistent dataset for constraining Cretaceous to mid-Tertiary exhumation histories (Fig. 4.7). Because many of the samples experienced slow cooling and resided in the partial annealing zone for extended periods of time (Fig. 4.7), the cooling paths modeled from the AFT data (using

HeFTy software; Ketcham, 2005; see Appendix B) tend to be more consistent and easily comparable than the ages themselves, so the models are more strongly emphasized in the following discussions. Additionally, because our AFT sampling mostly targeted samples above the PAZ, our results do not generally allow the depth range of the PAZ to be more precisely defined than by previous work by Colgan et al. (2004, 2006ab) and Stockli et al.

(2002). It is also important to note that the modeled late Tertiary reheating and final exhumation of the samples are constrained by model inputs based on geologic data

(Tables 4.1, 4.2; Fig. B3) and are not fully independent results. Although AFT-modeled results are shown in the range of 0–120 °C (Fig. 4.7), the technique is not very sensitive at temperatures below 60 °C. Near 120 °C, the models are also permissive of a wider range of cooling histories, frequently including older, more slowly cooled paths not supported by higher-T thermochronometry. AFT data from shallow levels (Du less than about 1 km) used in Colgan et al. (2006ab) were remodeled according to the methods 162

Figure 4.7. (Continued on next three pages) Thermochronology data, colored according to sample paleodepth (Du). A visual explanation is shown on the following page. Only “good” (Kolmogorov-Smirnov p > 0.5) AFT model results are shown. “Accept- able” fits (Kolmogorov-Smirnov p > 0.05) are shown for samples when no good results were produced (see Fig. B3 for “acceptable” results from other samples). Ages symbol- ized by rectangles show 2σ analytical uncertainty in both age and closure temperature. Colored lines labeled by depth represent best-fit paths to the thermochronology data from different depths, and represent the expected cooling paths taken by samples from the paleodepths shown. The temperature separation between these lines is drawn assuming a geothermal gradient that decreases linearly at 0.2 °C/km/m.y. to 27 °C/km at 16 Ma (except in the Santa Rosa Range; see text). 163 Age (Ma) 120 110 100 90 80 70 60 50 40 30 20 10 0 0 A. PINE FOREST 0 km RANGE 50 1 km

2 km 100 Temperature (°C) 3 km 150 4 km 0

5 km 1 200 2 Explanation 3 250

4 good AFT (km) u model results 300 D 5

acceptable AFT 6 350 model results 7

apatite U-Th/He 8 forward model 400 110 100 90 80 70 60 50 40 30 20 10 0 apatite U-Th/He 0 inverse model B. SANTA ROSA 0 km

zircon RANGE 50 1 km U-Th/He age 2 km K-feldspar MDD 100 Temperature (°C) model results ( 90% C.I. of median) 3 km (90% C.I. of distribution) 150

K-feldspar 4 km bulk age 200 5 km biotite 40 39 Ar/ Ar age 250

0 km best-fit thermal histories, as a 300 1 km function of Du 350 Colors indicate sample Du. All rectangles indicate 2σ 400 120 110 100 90 80 70 60 50 40 30 20 10 0 0 C. BLOODY RUN HILLS 50

100 0 km Temperature (°C)

1 km 150

2 km 200

250

300

350

400 Figure 4.7. (continued) 164

Age (Ma) 120 110 100 90 80 70 60 50 40 30 20 10 0 0 D. SEVEN TROUGHS RANGE 50 0 km

1 km 100

150

0 Temperature (°C)

1 200 2 3 250 4 (km) u 300

D 5 6 350 7 8 400

450

500

120 110 100 90 80 70 60 50 40 30 20 10 0 0 E. BLACK ROCK RANGE 50

100 Temperature (°C)

150 0 km

1 km 200

250

300

350

400

120 110 100 90 80 70 60 50 40 30 20 10 0 0 F. DIAMOND MOUNTAINS 0 km Temperature (°C) 1 km 50

100

150

200 Figure 4.7. (continued) 165

Age (Ma) 120 110 100 90 80 70 60 50 40 30 20 10 0 0 G. SELENITE RANGE Temperature (°C) 50

0 km 1 km 100

150

200 120 110 100 90 80 70 60 50 40 30 20 10 0 0 H. SAHWAVE RANGE 0 km 1 km 50

2 km 100 Temperature (°C) 3 km 150

200

250

300

350

400

120 110 100 90 80 70 60 50 40 30 20 10 0 0 I. WASSUK RANGE 50 0 km 100 1 km

2 km 150 Temperature (°C)

3 km 0 200 1 4 km

2 5 km 250 3 6 km 4 300 (km) u 7 km

D 5 350 6 8 km 7 400 8

450

500 Figure 4.7. (continued) 166 described in Appendix B in order to more directly compare them to the new data. In general, the thermal histories derived from inverse modeling of the AFT data from samples at different paleodepths within the same transect overlapped well when the geothermal gradient was taken into account (Fig. 4.8). These AFT thermal histories were also generally compatible with the low-temperature end of K-feldspar 40Ar/39Ar MDD inverse models (Fig. 4.7). Where both were available, biotite and K-feldspar 40Ar/39Ar thermal histories are broadly consistent with each other (Fig. 4.7). However, only a limited number of biotite 40Ar/39Ar ages were collected because the results were expected to reflect primarily post-magmatic cooling, and only four of six K-feldspar samples yielded usable thermal histories due to problems with the method used and the quality of the K-feldspar crystals analyzed (see Appendix D).

Apatite and zircon U-Th/He single grain ages (Tables 4.2, C1, C2), frequently disagree with each other and the other results at analytical uncertainty, and were of limited usefulness for constraining thermal histories (Fig. 4.7). For the apatite U-Th/He results, forward modeling in HeFTy indicates that most of this scatter (Table C2) could have resulted from a combination of very slow cooling through the partial retention zone

(probably 0.3–3 °C/m.y. for a majority of samples; Fig. 4.7) and natural variation in diffusion parameters (cf. Colgan et al., 2006b; Unruh et al., 2007; see Appendix C).

Apatite U-Th/He forward models in which the activation energies of the grains had to be slightly adjusted to provide an acceptable fit (see Appendix C), are shown in Figure 4.7 for comparison, but were not used to provide thermal history constraints. Similar problems may affect the zircon (U-Th)/He system, which is less thoroughly studied (cf.

Reiners et al., 2004), but even so, the presence of zircon U-Th/He grain ages younger 167

Figure 4.8. (Next page) Comparison of apatite fission-track length models from different paleodepths within the same transect. Model results were “normalized” for paleodepth by shifting them down in temperature according to their Du using a geothermal gradient decreasing at 0.2 °C/km/m.y. to 27 °C/km at 16 Ma. This should yield the temperature expected at the level of the unconformity. Samples that yielded no good results, and samples from paleodepths below 2.5 km are not included in these comparisons; e.g., “good fit in all models” means good fit in all samples above 2.5 km that yielded any good fits. 168 Time before present (millions of years) 100 90 80 70 60 50 40 30 20 10 0

20 A. PINE FOREST RANGE 60 Good fit in all models*

100 Acceptable fit in all models* Acceptable fit in a majority of models*

20 B. SANTA ROSA RANGE

100

20 C. BLOODY RUN HILLS

100

20 D. SEVEN TROUGHS RANGE

100

20 E. BLACK ROCK

Temperature (°C) Temperature RANGE 60

100

20 F. DIAMOND MOUNTAINS

100

20 G. SELENITE RANGE

100

20 H. SAHWAVE RANGE

100

Figure 4.8. (Continued) 169 than apatite fission track ages for some samples (Tables 4.2, C1) indicates that many of these data may simply be unreliable (see Appendix C), though in some cases they are still consistent with the other data (Fig. 4.7).

The various thermochronology data were used to construct visual best-fit thermal histories that were consistent between samples from different paleodepths (lines color- coded by Du, Fig. 4.7). In the following sections, each sample transect is addressed separately to provide geologic context for specific thermochronology results and interpretations. Emphasis is placed on the thermal record of Cretaceous to mid-Tertiary erosional exhumation rather than late Tertiary burial and extensional unroofing.

A. Pine Forest Range

The basement framework of the Pine Forest Range (Fig. 4.3) includes Paleozoic and early Mesozoic arc strata, deformed and metamorphosed in the early Jurassic and intruded by mid-Cretaceous plutons (Fig. 4.4A; Wyld, 1996). Plutonic rocks include a large, concentrically zoned, ~108 Ma intrusive complex spanning the northern part of the range (Fig. 4.4A; e.g., Wyld and Wright, 1997, 2001). The Tertiary unconformity is overlain by bimodal composition lavas of Oligocene age, which are substantially thicker in a paleo-valley mapped across the SW part of Fig. 4.4A (Colgan et al., 2006a, 2010).

These flows are in turn capped by ~16 Ma Steens basalt and younger Miocene ash-flow tuffs, forming a total Tertiary section about 1100 m thick (Fig 4.4A, 4.4B; Colgan et al.,

2006b). The basal Tertiary units dip about 30° to the west, and the range is bound by an east-dipping (measured at 40°) normal fault on its east side (Fig 4.4A, 4.4B; Colgan et al,

2006a). Existing thermochronology indicates that the top of the Miocene apatite fission 170 track PAZ begins about 700 m below the unconformity (Colgan et al., 2006ab).

The apatite He and fission track transect of Colgan et al. (2006a,b) is very thorough at greater paleodepths (more than 4 km), but does not target rocks closer to the

Tertiary unconformity as this study sought to avoid thermal disturbance from Eocene hypabyssal trachyandesite intrusions in the area (Fig. 4.4A; Colgan et al., 2006a).

Because of this, our transect includes a sample (PF-01) nearer the unconformity (Du ~

0.09 km) collected from a more northerly transect (Fig. 4.4A, 4.5A; Table 4.2). AFT- modeled cooling paths for this sample roughly overlap with those of the other shallow (<

0.6 km) samples of Colgan et al. (2006ab), particularly the younger ones that show moderately high cooling rates (~10 °C/m.y.) at ca. 77 Ma (Fig. 4.7A, 4.8). Cooling at this time would also be bracketed by the zircon U-Th/He grain ages for sample PF-01

(green boxes; Fig. 4.7A). Zircon U-Th/He grain ages from a much deeper sample (PF-

11; Du ~ 3.81 km; red boxes in Fig. 4.7A) show a considerable spread in age but coincide with very low cooling rates in the AFT-modeled cooling paths. The estimated temperature difference suggests a geothermal gradient on the order of 35 ± 15 °C/km at ca. 30 Ma, similar to the 12 Ma estimate of 27 ± 5 °C/km of Colgan et al. (2006b). K- feldspar MDD-modeling of sample PF11 was not successful due to experimental problems (see Appendix D), but by comparison to the successfully modeled K-feldspars, its bulk closure temperature is probably in the range of 210–490 °C (pink box, Fig. 4.7A).

B. Santa Rosa Range

In the Santa Rosa Range (Fig. 4.3), Upper Triassic basinal metasedimentary rocks are intruded by the 105.1 ± 1.1 Ma (Colgan et al., 2006b) Santa Rosa pluton and the 171 smaller Sawtooth stock (Fig. 4.4B; Compton, 1960), which is probably 85 Ma (Brown et al., 2010). The basement rocks are capped by 16.7 – 14 Ma volcanic rocks, which range from basalt to rhyolite and contain a greater proportion of intermediate composition lavas than found in surrounding areas (Fig. 4.4B; Brueseke and Hart, 2008, 2009). These volcanic rocks, about 800 m thick (LeMasurier, 1965), are titled eastward about 15°

(Compton, 1960) and the range is bound by a major west-dipping (35–40° dip) normal fault developed along the west side of the range (Fig 4.4B, 4.5B; Colgan et al., 2004).

Tilting exposes rocks up to around 4 km below the unconformity, but the distribution of

Cretaceous intrusions precludes collecting a continuous transect across the range (Fig.

4.5B). Existing thermochronology defines the top of the mid-Miocene apatite fission track PAZ to be at ~ 900 m below the unconformity (Colgan et al., 2004, 2006b).

Previous sampling above this level was limited (two AFT samples), partly because of the distance between the Santa Rosa pluton and the unconformity along the transect of

Colgan et al. (2004, 2006b). Our main transect across the east half of the range was chosen further south, where the Santa Rosa pluton is wider and the range crest is further east (Fig. 4.4B), such that the peak of the crest is within ~170 m of the projected unconformity (Fig. 4.5B).

AFT-modeled cooling paths and K-feldspar 40Ar/39Ar MDD modeling of a deep sample (SR-17; Du ~ 3.93 km; thick red curve in Fig. 4.7B) are generally consistent, showing moderate cooling rates (~2–10 °C/m.y.) until ca. 40 Ma followed by slower cooling (~0.1–1 °C/m.y.) until the deposition of overlying Miocene volcanic rocks (Figs.

4.7B, 4.8). The SR-17 MDD model results are also roughly consistent with rapid

Miocene exhumation by normal faulting (Fig. 4.7B), although this last part of the K- 172 feldspar cooling history by be compromised, since Miocene volcanic burial represents a small deviation from the assumption of monotonic cooling used in generating the model.

Best-fit forward models compatible with the apatite U-Th/He grain ages from two samples are broadly consistent with the above data. With the exception of three obviously spurious grain ages from sample SR-20 (Table C1), zircon U-Th/He ages are also consistent with the thermal history implied by the other data (Fig. 4.7B).

Taken at face value, the temperature differences among AFT and MDD model results from different paleodepths implies the presence of a very steep geothermal gradient, 40 ± 2 °C/km at 30 Ma (Fig. 4.7B; cf. 30 ± 6 °C/km at 12 Ma, Colgan et al.,

2006b). The apparent fanning of AFT and MDD model results from different Du at older ages (Fig. 4.7) also implies that the geothermal gradient was decreasing over time.

Similar results are noted in other transects and will be discussed in a later section, but the apparent geothermal gradient and its rate of change average about 30–40% higher in the

Santa Rosa Range than in other transects, including the nearby Bloody Run Hills (Fig.

4.7C; see below). Although this might reflect a local thermal anomaly, a simpler explanation would be that the Santa Rosa Range is more steeply tilted on its west side

(Fig. 4.5B) than dips extrapolated from the east side might suggest—an average tilt closer to 20° would result in calculated paleodepths about 35% greater and decrease the apparent geothermal gradient. Greater back-tilting of the range is also consistent with the low dip of the range-bounding normal fault (35–40°; Fig. 4.5B), since it is likely to have formed at close to 60° to the paleo-surface.

173

C. Bloody Run Hills

The Bloody Run Hills are effectively the southernmost part of the Santa Rosa

Range (Fig 4.3), and are treated as such in many older works and maps (e.g., Compton,

1960). They share similar metamorphic and igneous basement; basinal metasedimentary rocks are intruded by the 97.7 ± 0.7 Bloody Run Stock (Fig. 4.4C; Colgan et al., 2006b).

Both ranges are faulted on their western side and are titled to the east (Figs. 4.5B, 4.5C).

Tertiary strata are poorly exposed, but presumed to be similar in age and thickness to those in the Santa Rosa Range (Figs. 4.5B, 4.5C; Compton, 1960; Colgan et al., 2006b).

Colgan et al. (2006b) conservatively assume a 10° dip (not directly measurable) for the

Tertiary sediments. In fact, the proximity of Bloody Run Peak (Fig. 4.4C), which is granodiorite, to lower elevation Tertiary outcrops, suggests the unconformity (if roughly planar) must dip at least 13.5° to clear the top of the range (Fig. 4.5C). Analysis of the 3-

D outcrop pattern suggests a dip closer to 15°, the same as in the Santa Rosa Range, and a strike of N 35° E. Even with a 15° dip, Bloody Run Peak would be within ~100 m of the unconformity (shallower than our easternmost sample, Table 4.2), so our transect is focused on its western flank to gain more precise control on the relative depths of the samples, which vary from about 100 to 1800 meters below the unconformity (Figs. 4.4C,

4.5C).

AFT models indicate slow cooling (1–3 °C/m.y.) from ca. 75–20 Ma (Figs. 4.7C,

4.8). K-feldspar MDD modeling for sample BN-11 indicates more rapid cooling (up to

10 °C/m.y) from 90–75 Ma (thick blue curve in Fig. 4.7C). At its low temperature end,

40 39 the K-feldspar Ar/ Ar MDD modeled cooling history for sample BN-11 (Du = 1.73 km) overlaps with the AFT-modeled history from the same sample, providing a 174 continuous thermal history spanning 90 million years and 300 °C (Fig. 4.7C). Biotite

40Ar/39Ar results for sample BN-11 (empty blue box in Fig. 4.7C) also agree with the K- feldspar MDD results, but zircon U-Th/He results from the same sample (solid blue boxes in Fig. 4.7C) appear to be consistently too young. Three apatite U-Th/He single grain results for sample BN-01 were modeled together to yield the only acceptable inverse model of U-Th/He data in this study (hashed pattern in Figure 4.7C). Apatite U-

Th/He results from sample BN-05 could be fit to a reasonably consistent forward model

(teal line, outlined in black, Fig. 4.7C) by minor adjustments to the activation energies of the grains (See Appendix C). The geothermal gradient is not well constrained for this transect, because the range of paleodepths explored is not large, but averaged over 20-60

Ma, AFT model results are consistent with gradients in the range of 16 ± 13 °C/km.

D. Seven Troughs Range

Previous work in the Seven Troughs Range (Fig. 4.3) included only reconnaissance-level work, with focused studies in the Seven Troughs mining district along the east side of the range (e.g., Ransome, 1909; Hudson et al., 2006). The geologic map presented in Fig 4.4D represents new reconnaissance mapping integrated with the

1:250,000 county map by Johnson (1977). The Seven Troughs Range is underlain by slates and phyllites of the basinal terrane, of presumed Late Triassic/Early Jurassic age

(Fig. 4.2, 4.4D; e.g., Ransome, 1909; Johnson, 1977). In the northwestern Seven

Troughs Range (area of Fig. 4.4D), these rocks are intruded by a Cretaceous tonalite whose zircons yielded a 207Pb corrected 206Pb/238U weighted average age of 102.9 ± 1.5

Ma (2σ error; Fig. 4.6; see Appendix E for SHRIMP analytical data). Ash-flow tuffs are 175 present at the base of the overlying Tertiary section in many places (though not on the line of section; Fig. 4.5D). In the southeastern Seven Troughs Range these undated basal tuffs have been tentatively correlated with the ca. 30-23 Ma ash-flow tuffs that fill paleotopography in nearby ranges (Hudson et al., 2006). Most of the Tertiary units are

14.4 – 13.8 Ma bimodal volcanic rocks that fill a syn-magmatic graben system in the eastern Seven Troughs Range (SE corner of Fig. 4.4D; Hudson et al., 2006). This graben system also hosts a number of hypabyssal rhyolite and basalt intrusions (SE corner of

Fig. 4.4D) related to epithermal gold-silver mineralization (Hudson et al. 2006). Outside the graben, Tertiary strata reach an exposed thickness of ~300 m beneath the peaks of the range, but this only represents a minimum estimate of the original thickness, as post-14-

Ma erosion in this area is expected to be substantial. Although Tertiary strata are flat lying in the southeast Seven Troughs Range (Hudson et al., 2006), they dip gently to the northeast in the area of Fig. 4.4D, estimated at ~8° along the line of section, exposing rocks up to about a kilometer deep beneath the unconformity (Fig. 4.5D). Because basement rocks are exposed along the western side of the range, a west-dipping normal fault is inferred to bound the range on this side (Fig. 4.4D, 4.5D), although it is not clearly expressed by topography (Fig. 4.3).

About 30 m below the unconformity, sample 7T-1A yielded an apatite-fission- track age of 16.5 Ma, which presumably reflects thermal resetting by a thick rhyolite flow that overlies this part of the unconformity. Deeper AFT samples reveal steady cooling at a rate of ~ 1.5 °C/m.y. from 74–30 Ma (Figs. 4.7D, 4.8). This part of the cooling history can be matched by forward modeling of eight apatite U-Th/He grain ages from two

40 39 samples (Fig. 4.7D). K-feldspar Ar/ Ar MDD modeling of sample 7T-5 (Du = 0.79 176 km) suggests rapid cooling (~100 °C/m.y.) at 77–74 Ma (Fig. 4.7D). Biotite 40Ar/39Ar and zircon U-Th/He ages from the same sample, as well as AFT model results, are fairly consistent with this K-feldspar-MDD-modeled thermal history (Fig. 4.7D). If the implied rapid cooling is not just a spurious result, it could have been achieved either by significant tectonic exhumation (at rates around 3 mm/a) or by intrusive reheating.

Although a few authors have speculated on tectonic deformation in this general time and area (e.g., Moores et al., 2002), structures active at this time in NW Nevada have never been identified during detailed mapping (e.g., Compton, 1960; Wyld, 1996; Van Buer, submitted). The simplest explanation is probably thermal resetting by a hypothetical nearby intrusion at 76 Ma, which appears as a plateau in the age spectrum (Table D2; Fig.

D2; but note that the K-feldspar 40Ar/39Ar total gas age is 86.2 ± 1.0 Ma). U-Pb dating has not revealed any intrusions of this age in the region, however (e.g., Van Buer and

Miller, 2010).

E. Black Rock Range

The basement framework of the Black Rock Range is similar to that of the Pine

Forest Range to the north (Fig. 4.2, 4.3), with Permo-Triassic arc strata extensively intruded by Cretaceous plutons (Fig. 4.4E; e.g., Lerch et al., 2008b). Near our transect

(Fig. 4.4E) the granodiorite has been dated at 112.3 ± 1.4 Ma (Colgan et al., 2006b).

Basement rocks are overlain unconformably by 35 Ma basalt and Oligocene bimodal volcanic rocks for a total Tertiary thickness of ~1 km (Fig. 4.4E, 4.5E; Colgan et al.,

2006b; Lerch et al., 2008). Miocene ignimbrites are present in the Black Rock Range, but not in the area of Fig. 4.4E. The Black Rock Range is faulted along both sides, 177 although at the latitude of Fig. 4.4E, the east-side faults dominate, exposing basement and tilting Tertiary strata gently to the west (Fig. 4.5E; Lerch et al., 2008b). The presence of multiple range-front faults (Fig 4.4E) limits sampling to a ~350 m maximum paleodepth

(Fig. 4.5E).

AFT-modeled cooling paths from different samples mostly overlap very well

(Figs. 4.7E, 4.8), except for JC02-BL5, which is almost 10 Ma older (Table 4.2), and for which only 12 grains were counted for age determination (Colgan et al., 2006b). In most of the AFT models, moderately fast cooling (~8 °C/m.y.) before 75 Ma is followed by slow cooling (~0.5 °C/m.y.) until latest Eocene deposition and burial by volcanic rocks

(Fig. 4.7E). The biotite 40Ar/39Ar age of 105.3 ± 1.3 Ma probably records the last stages of post-magmatic cooling, but is almost on-trend with cooling rates extrapolated from the younger AFT data at low temperature, which should reflect only erosional exhumation

(Fig. 4.7E). This suggests that most of the intervening cooling (Fig. 4.7E) was related to gradual erosional exhumation.

F. Diamond Mountains

The Diamond Mountains are generally considered to be the northernmost part of the Sierra Nevada mountain range (Fig. 4.3), but have received comparatively little study relative to other parts of the range. Basement rocks are almost entirely Cretaceous batholithic rocks (Fig. 4.4F; Oldenburg, 1995). Zircon U-Pb ages from sample DD-8 on our transect are spread out in a bimodal distribution spanning 103 – 109 Ma, likely due to recycling of zircon from older intrusions (Fig. 4.6). The younger peak of this distribution gives an age of 104.5 ± 1.3 Ma (2σ error; Fig. 4.6; see Appendix E for SHRIMP data). In 178 the area of Fig. 4.4F, the oldest Tertiary rocks are the 16 Ma Lower Basalt of Thompson

Peak, presumed to correlate with the Lovejoy Basalt (Grose and Porro, 1989; Roberts,

1985). This basalt is not exposed continuously; in many areas volcaniclastic gravels or younger andesite and basalt flows (as young as 10 Ma) overlie the unconformity directly

(Fig. 4.4F). In other parts of the Diamond Mountains, though not within the area of Fig.

4.4F, the bedrock beneath the unconformity is cut by paleochannels containing gravels

(e.g., Grose and Porro, 1989). These were once thought to be Lower Eocene by comparison of fossil plants to the auriferous gravels elsewhere in the Sierra Nevada

(Wolfe, 1994), but are more likely Upper Eocene given the presence of 35 Ma detrital zircon (Lovelock, 2010). Elsewhere in the range, Oligocene ignimbrites also fill these paleochannels (Hinz et al., 2009). Tertiary strata have a demonstrated thickness of ~300 m beneath Thompson Peak (Fig. 4.4F), but this is clearly just a lower bound on the pre- extensional thickness of these units.

The Diamond Mountains have been uplifted along their northeastern edge by the

Honey Lake Fault Zone (Fig. 4.4F), which marks the northernmost extent of the Sierra

Nevada frontal fault system (Fig. 4.3), and frequently consists of several subparallel faults (e.g., Roberts, 1985). The uplifted Tertiary strata remain more or less flat-lying

(Fig. 4.5F; Grose and Porro, 1989). Our transect crosses the range front south of

Susanville (Fig. 4.3), at the location of greatest relief within this fault block (Fig 4.5F).

Here the Tertiary unconformity lies ~800 m topographically higher than the lowest outcrops of basement rocks (Fig. 4.5F).

Samples from the Diamond Mountains produced a low yield of apatite of mixed quality; many grains demonstrated strong zoning and high concentrations of lattice 179 dislocations. Sample DD-1 did not have enough measurable confined fission tracks to model track length distributions; its age is shown with the green box in Fig. 4.7F. AFT- modeled cooling paths for the other two samples demonstrate moderate cooling rates (~3

°C/m.y.) until ca. 60 Ma, followed by very little further cooling after this time until

Miocene extension (Figs. 4.7F, 4.8). Best-fit forward models for 7 single-grain apatite U-

Th/He ages from two samples are broadly consistent, but suggest that more prolonged slow cooling may be needed to fit the younger grain ages (Fig. 4.7F). However, as high-

U rims were observed on many apatites during AFT work, anomalously young U-Th/He ages may be suspect (see Appendix C).

G. Selenite Range

The northern Selenite Range (Fig 4.3) has been mapped at 1:250,000 scale by

Johnson (1977), but the geologic map in Fig. 4.4G is drawn primarily from reconnaissance mapping during this study. The basement on the west side of the range is composed of Permian metavolcanic rocks and lower Mesozoic metasedimentary rocks, which include phyllite, marble, and lenses of gypsum (Fig. 4.4G; Johnson, 1977). These are intruded to the east by Cretaceous granodiorite that has been dated by the U-Pb method on zircon at 96.4 ± 1.2 Ma (2σ error; Fig. 4.6; see Appendix E for SHRIMP analytical data). Tertiary strata in the area are primarily basalt flows (Fig. 4.4G) that remain undated, but are perhaps 18–14 Ma based on correlation to similar units mapped and dated in the southernmost Selenite Range by Whitehill (2009). On the west side of the range in this area, Tertiary strata are mainly confined to a north-south-trending paleochannel (Fig. 4.4G), which is floored by discontinuous outcrops of petrified-wood- 180 bearing, basement-clast conglomerate locally overlain by ash-flow tuffs of unknown age.

Unfortunately, these Tertiary outcrops are separated from most of the granodiorite by a west-dipping normal fault internal to the range (Fig. 4.4G), making them useless as a datum for paleodepth. On the east side of the range, a minimum 200 m thick section of basalt is downdropped by an east-dipping normal fault (Fig 4.4G, 4.5G). These strata dip

~11° to the northeast, almost parallel to the strike of the fault, suggesting that this block, which trends obliquely to the range, is part of a large lateral ramp in the normal fault system that bounds the central part of the east side of the range (Fig. 4.4G). This fault block provides basement exposure up to about 700 m structurally below the unconformity

(Fig. 4.5G).

Modeled cooling histories for two AFT samples from the Selenite Range agree well with each other, with cooling rates slowing from ~4 °C/m.y. to < 1 °C/m.y. at about

50 Ma (Fig. 4.7G, 4.8). Forward modeling of four single-grain apatite U-Th/He ages matches this pattern as well.

H. Sahwave Range

The basement of the Sahwave Range (Fig. 4.3) is dominated by the 93–88.5 Ma

Sahwave Intrusive Suite, a large (~1000 km2), concentrically zoned intrusion (Fig. 4.4H;

Van Buer and Miller, 2010). Section HH´ (Fig. 4.4H) is within the ca. 93 Ma

Granodiorite of Juniper Pass, the outermost and oldest unit of the suite (Van Buer and

Miller, 2010). This intrudes lower Mesozoic metasediments of the basinal terrane (Fig.

4.2; Van Buer and Miller, 2010), but in the area of Fig. 4.4H these are only present as stoped blocks or roof pendants tens of meters to a few kilometers in scale (Plates 7, 8). 181

Tertiary strata have been eroded away from most of the range, but small exposures can be found around the edges (Fig. 4.4H). 31–25 Ma Oligocene ash-flow tuff fills a paleovalley in the southern Nightingale Range (Fig. 4.3) just southwest of Fig. 4.4H

(Plate 6; Faulds et al., 2005), but is not found within the Sahwave Range itself. Limited outcrops of basalt, dated at 13.3 ± 0.8 Ma (Whitehill, 2009), mark the only exposures of the Tertiary unconformity around the western Sahwave Range (Fig 4.4H; Plate 7).

Within the main Sahwave fault block, these basalts dip ~14° to the west, but their dips decrease westward towards the axis of a syncline between the Sahwave Range and the

Nightingale Range to the west (Fig 4.5H; Plate 7; Whitehill, 2009; Van Buer, submitted).

The thickness of pre-extensional Tertiary strata is constrained to be at minimum ~300 m from correlative outcrops in the northern Nightingale Range (Fig. 4.3; Plates 3, 4), and at maximum ~800 m from un-reset apatite U/Th-He ages from Whitehill’s (2009) thermochronology transect just north of Fig. 4.4H. The east side of the Sahwave Range is marked by two generations of normal faults: an older, low angle normal fault, back- rotated to ~32° (measured in outcrop), and an active, high angle fault, which steps eastward into the basin, uplifting Tertiary basin-fill in its footwall (Figs. 4.4H, 4.5H;

Plate 8; Whitehill, 2009; Van Buer, submitted). This Tertiary package is dominated by fanglomerates, sandstones, and lacustrine deposits, and demonstrates progressive tilting from ~28° at its base to ~11° at its top, and is dated at ca. 11 – 9 Ma where constrained by ages of interlayered basalt and tephra (Fig. 4.4H, 4.5H; Plate 8; Whitehill, 2009).

Given the low angle of the older normal fault and the up to ~28° dip of its hanging-wall sediments, the projected unconformity is expected to steepen towards the fault, as observed in better-constrained large-offset normal faults (e.g., Egger and Miller, 182

2011; Fig. 4.5H; Plates 7, 8). Our cross section, which assumes the projected unconformity reaches a 28° dip where it intersects the fault and minimizes its curvature, suggests paleodepths up to ~2800 m are exposed below the unconformity (Fig. 4.5H).

Previous thermochronology from a transect just north of the area of Fig. 4.4H was inferred to reach greater paleodepths (Whitehill, 2009; Plate 5), but structural control is limited and new mapping suggests the transect may cross a substantial normal fault near its west end (Van Buer, submitted; Plate 4). The transect for this study was chosen to provide tighter control on paleodepths near the unconformity, which has limited exposure in the Sahwave fault-block (Figs. 4.4H, 4.5H; Plate 7).

Fission track data define the top of the pre-extensional partial annealing zone to be at a paleodepth (Du) of 1.9 ± 0.3 km (Fig. 4.9). Assuming 0.5 ± 0.3 km of overlying

Tertiary strata, surface temperatures of 10 ± 5 °C, and that the top of the PAZ represents a temperature of 60 ± 5 °C, this suggests a Miocene geothermal gradient of roughly 21 ±

5 °C/km, which is also consistent with the temperature spread of AFT model results from different paleodepths (but note that Whitehill, 2009, estimated ~37 °C/km nearby). AFT- modeled cooling paths for the samples suggest steady cooling at 1–2 °C/m.y. from 80 Ma until Miocene volcanic burial (Figs. 4.7H, 4.8). Extended residence at ~68 °C is needed to fit the spread in apatite U-Th/He grain ages from sample SE-04, but these ages may not be entirely reliable (see Appendix C). Zircon U-Th/He grain ages from this transect are too scattered to give a consistent interpretation (Table C1, Fig. 4.7H) K-feldspar

40 39 Ar/ Ar step-heating results from sample SE-12 (Du ~ 2.76 km) featured an intermediate age maximum (Fig. D2) and could not be adequately modeled (see

Appendix D). However, the total gas age for this K-feldspar sample is 81.0 ± 0.4 Ma, 183

Apparent fission-track age (Ma) 80 70 60 50 40 30 20 10 0 H. SAHWAVE RANGE maximum thickness Tertiary strata minimum thickness Tertiary strata

0 SE-01 Tertiary unconformity SE-04 in km) u

1 SE-06

SE-08

Paleodepth (D ~60 °C 2 SE-10 exhumed Miocene AFT PAZ SE-12 3

Figure 9. Apatite fission-track ages from the Sahwave Range (2σ errors) plotted against paleodepth suggest the top of the AFT partial annealing zone lies between 1.6 - 2.2 km below the unconformity. 184 indicating that the sample was probably cooling fairly rapidly until 80 Ma, after which the AFT data indicate slower cooling (Fig. 4.7H). This rapid cooling presumably reflects the final stage of post-magmatic cooling of the Sahwave Intrusive Suite.

I. Wassuk Range

The basement of the northern Wassuk Range (Fig. 4.3) is composed of lower

Mesozoic metavolcanic rocks intruded by Cretaceous quartz monzonite (Fig. 4.4I;

Bingler, 1978; Surpless, 1999; Stockli et al., 2002). The quartz monzonite has a biotite

K/Ar age of 80 ± 3 Ma (Bingler, 1972), but crystallization ages are usually at least ~5 Ma older than biotite Ar ages on this side of the Sierra Nevada batholith (e.g., Stern et al.,

1981), and a similar unit in the southern Wassuk Range gives a zircon U-Pb age of about

88 Ma (Klinger, 2005). Basement units are affected by sodic hydrothermal alteration

(Surpless, 1999), and consequently our apatite separates from Wassuk Range samples contained consistently low quality, fluid-inclusion-riddled grains. The oldest dated

Tertiary units above the unconformity are 27–26 Ma ash-flow tuffs, which fill paleotopography and occasionally overly a thin, undated boulder conglomerate (Fig. 4.4I; e.g., Surpless, 1999). These tuffs and overlying 15–14.5 Ma andesite flows predate extensional faulting and are tilted about 60° to the west (Fig. 4.4I; 4.5I; Surpless, 1999).

The total thickness of these units varies according to the underlying topography, but is estimated to be about 800 m near the line of section (Fig 4.5I; Surpless, 1999; Stockli et al., 2002). Younger 14.5–7 Ma sediments and basalts, dipping only 17–43°, are deposited across an angular unconformity above older Tertiary strata and basement rocks

(Fig. 4.5I; Surpless, 1999). As in the nearby Yerington district (Proffett and Dilles, 185

1984), the northern Wassuk Range is affected by three generations of east-dipping normal faults accommodating nearly 300% extension, with the oldest faults back-rotated almost to horizontal, and the youngest faults bounding the modern range at a high angle

(Fig. 4.4I, 4.5I; Surpless, 1999; Stockli et al., 2002; Surpless et al., 2002). This faulting has exposed basement rocks that resided at paleodepths greater than 7 km structurally beneath the unconformity (Fig. 4.5I). Our sample transect is located north of previous thermochronological work (Surpless, 1999; Stockli et al., 2002; Surpless et al., 2002) to take advantage of granitic outcrops closer to the unconformity, and is arranged more nearly perpendicular to the strike of the unconformity and range-front fault (resulting in small reductions to the paleodepth estimates of Surpless, 1999, and Stockli et al., 2002;

Fig. 4.4I, 4.5I). Previous thermochronology defines the top of the apatite fission track and U-Th/He partial annealing/retention zones at about 1400 m and 800 m below the unconformity, respectively (Stockli et al., 2002).

Unfortunately, the low quality of apatite grains from our northern transect, combined with the small number of grains to choose from (many samples yielded no apatite) enabled us only to measure one dubious nine-grain fission-track age of 124 ± 20

Ma (shown as the green box in Fig. 4.7I) and no track-length distributions (Tables 4.2,

B1). Figure 4.7I does show AFT modeling results from Surpless (1999) for sample

96BS-11.5a (Du ~ 0.8 km), which shows relatively slow, steady cooling from 70–40 Ma, and sample 96BS-11.4 (Du ~ 4.6 km), which records both Miocene and Pliocene

40 39 extensional exhumation. K-feldspar Ar/ Ar MDD modeling of sample WK-15 (Du ~

7.2 km) also records rapid Miocene exhumation beginning at ~ 16 Ma, preceded by more gradual cooling (Figure 4.7I). However, the deep K-feldspar-derived thermal history 186 demonstrates faster cooling than the AFT-modeled history from near the unconformity

(Fig. 4.7I), suggesting a decrease in geothermal gradient from about 28 °C/km at 35 Ma to about 22 °C/km at 16 Ma. Both estimates are within error of the 27 ± 5 °C/km estimate at 15 Ma of Stockli et al. (2002). Zircon U-Th/He grain ages are scattered, but not completely inconsistent with the other data (Fig. 4.7I).

INTERPRETATION OF THE THERMAL HISTORIES

Geothermal Gradient

Understanding the geothermal gradient and how it may have changed over time is essential for estimating the amount of exhumation represented by the thermal histories summarized in Fig. 4.7. Unfortunately, most of the transects studied lack sufficient variation in paleodepth to constrain the geothermal gradient precisely, so it is necessary to assume the geothermal gradient shared a similar history throughout the segment of the batholith studied. Published estimates of the Miocene geothermal gradient throughout this region actually cluster quite closely, in the range of 26–30 °C/km (Stockli et al.,

2002.; Colgan, 2006b). Although geothermal gradients estimated for our transects range from 16–40 °C/km, these results are generally consistent, within the estimated uncertainties, with geothermal gradients in the range of 26–30 °C/km. One exception is the abnormally high gradient (40 ± 2 °C/m.y.) estimated for the Santa Rosa Range, but this result may be an artifact of underestimated tilt of the range.

The geothermal gradient is most thoroughly constrained in the two transects with

40 39 deep samples (Du > 3 km) that yielded K-feldspar Ar/ Ar MDD-modeled thermal histories that overlapped in time with most of the AFT-modeled cooling paths: the Santa 187

Rosa Range and the Wassuk Range (Fig. 4.7B, 4.7I). In both cases, the pre-16-Ma geothermal gradient appears to decrease over time by around 0.2–0.3 °C/km/m.y. In both cases, more rapid changes in the geothermal gradient may be needed before ~38 Ma, assuming the K-feldspar results are reliable (Fig. 4.7B, 4.7I). In the case of the Santa

Rosa Range, if the geothermal gradient has been overestimated due to underestimation of paleodepths, its rate of change will be overestimated, too. Data from other sections do not require a decreasing geothermal gradient, but seem to be compatible with it (Fig. 4.7).

A decreasing geothermal gradient in the late Cretaceous to early Tertiary is compatible with cessation of arc magmatism followed by shallow-angle subduction, and changes in the geothermal gradient of similar timing and magnitude have been suggested by previous work in the western Cordillera (Dumitru et al., 1991). Thermal modeling of subduction refrigeration suggests that the rate of change of the geothermal gradient should itself decrease with time (Dumitru, 1990). This seems consistent with the greater divergence between K-feldspar and AFT thermal histories from the Santa Rosa and

Wassuk Ranges before ~ 38 Ma (Fig. 4.7B, 4.7I). However, it is difficult to reliably extrapolate these fluctuations in the geothermal gradient to other sample transects, especially for times earlier than constrained by the K-feldspar data (pre-55 Ma; Fig.

4.7B). Therefore, all thermal history data (except data from the Santa Rosa Range) were modeled using a geothermal gradient that decreased linearly at 0.2 °C/km/m.y., reaching

27 °C/km at 16 Ma, when constructing the best fit cooling paths for rock at different Du

(depth-labeled colored lines in Fig. 4.7) and estimating depth of exhumation. Cooling data from the Santa Rosa Range were better fit using geothermal gradients (and their rate of change) that are 35% higher. 188

Figure 4.10A shows the depth of the final Tertiary unconformity surface beneath the paleo-surface as a function of time for each of the nine transects, estimated from the best-fit thermal histories of Fig. 4.7 by assuming the changing geothermal gradient described above. In short, Figure 4.10A shows the progress of exhumation for the nine transects. The depth estimates are only valid in so far as (1) the assumption of a stable geotherm is reasonable, and (2) the assumed geothermal gradients are correct. For example, the very rapid periods of cooling recorded in the Seven Troughs and Sahwave

Ranges may represent post-magmatic cooling when thermal gradients were neither close to the modeled values nor at steady state (Figs. 4.7D, 4.7H, 4.10A). In general, the geothermal gradient is less well constrained at earlier times, so the oldest depth estimates are probably the least reliable (Fig. 4.10A). In particular, depth estimates before ca. 40

Ma may be overestimates if the geothermal gradient was as high as suggested by the K- feldspar MDD results of the Santa Rosa and Wassuk Ranges (Figs. 4.7B, 4.7I, 4.10A).

Pre-40 Ma depth estimates should probably be considered to be maxima. The total exhumation since intrusion is difficult to infer from these data because it is difficult to determine with certainty what part of the early cooling history was related to post- magmatic cooling versus exhumation. The reasonably steady cooling paths observed for the Bloody Run Hills and Black Rock Range transects (Fig. 4.7C, 4.7E) provide tentative evidence that the total Late Cretaceous to mid-Tertiary exhumation in these areas may have been around 6 km (Fig. 4.10A). This value is consistent with previous estimates of

3–8 km of denudation near the axis of the NW Nevada batholith (Colgan et al., 2006b;

Van Buer et al., 2009), suggesting that our assumptions about changes in the geothermal gradient are probably approximately correct. 189

Time before present (millions of years) 110 100 90 80 70 60 50 40 30 20 10 0 A.

2 WK 3 DD SL Average age of last 2.3 km exhumation (Ma) SR 80 70 60 50 4 80 PF B.

cooler WK 5 90

6 SE

e r BN e SE SL Modeled depth (km) r BN he e 7 h 100 sp BR osp h o t h 7T li it l DD SR 8 m m

k Intrusive age (Ma) k PF 0 0 hotter 110 5 6 9 g g BR n n i i l l

o o

o o 7T c c 10 Figure 4.10. (A.) Black curves compare the exhumation histories for each transect. Depths are calculated from the best-fit cooling histories shown in Fig. 4.7 using a linearly-decreasing geothermal gradient. In gray, the decrease in average temperature for lithosphere of 50 and 60 km thickness is shown, assuming cooling from beneath begins at 109 and 99 Ma, respectively (see text). This is plotted using the same time scale, with temperature in arbitrary units, decreasing upwards, as the y-axis. DD - Diamond Moun- tains, BR - Black Rock Range, PF - Pine Forest Range, SR - Santa Rosa Range, 7T - Seven Troughs Range, SL - Selenite Range, WK - Wassuk Range, SE - Sahwave Range, BN - Bloody Run Hills. (B.) The inset at lower right shows the correlation between the intrusive age (estimated 2σ error shown), which decreases to the east, and the average age of the last 2.3 km of exhumation (calculated by integration of the curves in panel A). 190

Variations in Exhumation among the Transects

Comparison of the exhumation paths between the different transects (Fig. 4.10A) supports the earlier result of Colgan et al. (2006b) that exhumation continued until later in the east than in the west, and extends this generalization over the broader area of the present study. Interestingly, this west-to-east progression of erosional exhumation seems to follow the same pattern as the west-to-east architecture of the arc itself. For comparison to previously published data, Figure 4.11 shows the average fission-track ages from within 600 m of the Tertiary unconformity from the present study and nearby areas. The line between shallow fission-track ages older and younger than 70 Ma roughly parallels the division between the Early and Late Cretaceous sides of the Sierra

Nevada Batholith (Fig. 4.11). The age of the local granitic basement is fairly well correlated with the average age of the last ~2 km of exhumation (Fig. 4.10B), as calculated from the best-fit thermal histories shown in Fig. 4.10A. Furthermore, the slope < 1 for this correlation suggests that exhumation of the older, western intrusions continued for a shorter duration since intrusion, about 30–40 million years, than the exhumation of the younger intrusions, which lasted closer to 50 million years past intrusion (Fig. 4.10B). A more subtle trend observable in Fig. 4.10A is that the late- exhumed, eastern transects underwent their final 2-3 km of exhumation at a relatively steady rate (~0.05 mm/a), whereas the early-exhumed, western transects display more asymptotic exhumation paths, with more rapidly decreasing exhumation rates. Although exhumational history is fairly well correlated with east-west variations in the arc, consistent north-south variations were not discerned (Fig. 4.11).

The observed correlations between exhumational history and the structure of the 191

Explanation of Map Units OR 42° OR 73 Average fission track ages 73CA N NV 75 Cenozoic cover 77

99-85 Ma intrusions 69 80 120-100 Ma intrusions 61 older intrusions and 58 intrusions of unknown age 41° N Other pre-Cenozoic rocks Geologic map after Van Buer and Miller (2010). 61 65 70 Ma CA

th 87 li o th a B 40° 64 N

ada

e i

S 60 Ma

70* 60*

72*

s u

SIERRA NEVADA o

39° ce a

N t 52

re 68

C 54

rl 55

a E

121° W 120° W 119° W 118° W Figure 4.11. Geologic map showing average apatite fission track age for samples less than 0.6 km below the Tertiary unconformity at several sites in the study area. Average38º ages including data from this study are shown in bold. Starred ages actually represent times of cooling below ~90 °C estimated fromYosemite the apatite and zircon U-Th/He results of Cecil et al. (2006). Other fission-track ages areVillage from Surpless et al (2002) and Colgan et al. (2006b). As delineated by the stylized age contours, ages older than 70 Ma are found in the western, Early Cretaceous part of the batholith, and 70 - 60 Ma ages are found in the eastern, Late Cretaceous part of the batholith (note that few rocks of this age extend north of 41°) . Somewhat younger ages are found in the area east of the main batholith, where the density of intrusions is lower. 192 batholith (Fig. 4.10B, 4.11) suggest a potentially causal relationship: exhumation might be driven by the decaying thermal anomaly of the magmatic arc. The dashed gray curves superimposed on Fig. 4.10A (using the same time scale, temperature increasing downwards) show the average temperature decrease over time in a lithosphere of 50 or 60 km thickness after cooler material is introduced (or a heat source is removed). The thermal approximation shown (Fig. 4.10A) is an analytical solution to the 1-D heat equation, assuming a homogeneous lithosphere with a thermal diffusivity of 1 mm2/s, no heat production, and a step-function decrease in temperature below the lithosphere at t =

0, relative to an initially linear geothermal gradient. Average temperature is displayed in arbitrary units relative to the initial change in temperature. The observed delay between intrusion and the cessation of exhumation is of similar duration as the cooling time for

50–60 km lithosphere (Fig. 4.10A; cf. Dumitru, 1990). Note that in this case, cooling begins quickly after intrusion, and might be explained by removal of the arc’s heat source without requiring immediate interaction with a shallow slab. The longer delay between intrusion and the end of exhumation in the younger, eastern transects (Fig. 4.10B) would be consistent with the presence of thicker lithosphere to the east. Reconstructed Late

Cretaceous crustal thicknesses in NW Nevada, based on seismic data and estimates of extension, magmatism, and erosion, also thicken to the east, increasing from ~34 to ~40 km, consistent with the tectonic setting along the western margin of the continent (Lerch et al., 2007).

It is also interesting to note the similarity in form between the exhumation and arc-cooling curves (Fig. 4.10A), which also suggests a connection between the cooling and exhumational histories. The average temperature of the lithosphere might be 193 expected to correlate with the exhumation rate rather than net exhumation, but the approximately exponential form of these curves (Fig. 4.10A) makes this a minor distinction. Alternatively, the rate of change of temperature should also be proportional to the geothermal gradient, which would also be expected to decrease exponentially with time. Because the depths shown in Figure 4.10A were modeled using a linear approximation to the decreasing geothermal gradient, it is possible that exhumation actually proceeded at a relatively constant rate, and much of the observed curvature is caused by exponential decay of the geothermal gradient.

DISCUSSION

Late Cretaceous to Mid-Tertiary Paleogeography

The apparent correlation between exhumation of the northernmost Sierra Nevada batholith and cooling of the arc lithosphere provides insight into the Late Cretaceous to mid-Tertiary paleogeography of the region. Figure 4.12 schematically shows the remaining exhumation and the rate of exhumation as a function of transverse distance across the arc (as inferred from the exhumation paths shown in Fig. 4.10) at four times between 90–30 Ma. This study only constrains exhumation at 90 Ma along the western edge of the arc, because the arc was still active along its eastern edge until ca. 85 Ma.

However, in analogy to modern arcs, the active portion of the arc is presumed to have been relatively high-standing, resulting in exhumation rates as least as high as in the extinct western part of the arc (Fig. 4.12). This study does not constrain the magnitude or rate of exhumation east of the batholith at any time, but an eastward decline in metamorphism from greenschist grade near the batholith to little to no metamorphism in 194

0 West Cross-arc distance East

1 A. 30 Ma

2 50 Ma 3 70 Ma 4

6 90 Ma 7 Remaining exhumation (km) higher B.

~ 50 km Topographic relief 90 Ma

0.2 70 Ma

0.1 50 Ma lower 30 Ma

Exhumation rate (mm/a) 0 110 100 90 Approximate average basement intrusive age (Ma) Figure 4.12. (A.) Schematic drawing of remaining exhumation as a function of east- west distance across the arc, extrapolated from the exhumation curves shown in Fig. 4.10. This diagram can also be thought of as a cross section showing the depth of rocks now at the Tertiary unconformity at the times shown. East-west distance also correlates well with the age of the batholithic rocks, as shown on the scale at the bottom of panel B. Dashed lines indicate extrapolation based on inference rather than thermochronological data. (B.) Schematic drawing of estimated exhumation rates across the arc. This probably approximates cross sections of topographic relief (but not necessarily elevation) at the times shown. 195 the Mesozoic synclines mapped beneath the Tertiary unconformity across central Nevada implies that the magnitude of exhumation generally decreases away from the axis of the batholith (Van Buer et al., 2009). This conclusion is consistent with thermally mediated uplift and erosion of the arc. In general, exhumation rates decreased substantially over time in all parts of the arc after magmatism ceased, but the rate of exhumation decreased most rapidly to the west, causing the area of fastest exhumation to move gradually eastward with time (Fig. 4.12).

Erosional exhumation focused on the arc is presumably linked to the presence of topography that is higher or has greater relief than surrounding areas. Therefore, the rate of exhumation across the arc probably approximates its relative topographic relief (Fig.

4.12B). This supports the idea that the arc was a relatively high-standing feature when it was active, and may have acted as a drainage divide in the Late Cretaceous (e.g.

Wakabayashi and Sawyer, 2001; DeGraaff-Surpless et al., 2002, Van Buer et al., 2009).

Shortly after peak intrusion rates around 90 Ma (e.g., Ducea, 2001; Van Buer and Miller,

2010), the exhumation rate along the axis of the batholith began to decrease roughly exponentially, halving every ~25 million years, so the topographic relief may have decreased similarly (Fig. 4.12B).

Consistent with the pattern of exhumation rates extrapolated from our data, which decrease especially rapidly along the western side of the batholith (Figure 4.12), deposition of auriferous gravels across the western Sierra Nevada also indicates that exhumation had mostly ceased along the western edge of the batholith by ~50 Ma. Note, however, that the Eocene gravels rest only on Early Cretaceous batholithic rocks and older units, and only in the deepest of the paleochannels (Cassel and Graham, 2011). 196

Large rounded boulders also indicate that significant relief still existed nearby to the east

(Cassel and Graham, 2011), and abundant detrital zircons from the ca. 90 Ma intrusive suites of the eastern Sierra Nevada Batholith indicate that exhumation was still ongoing across the study area as these sediments were being deposited (Cassel et al., in press). By

~35 Ma, volcanic units were deposited above the unconformity in the Pine Forest and

Black Rock Ranges, in the Early Cretaceous part of the arc (Lerch et al., 2008a). Also at this time, 55–33 Ma detrital zircon grains, apparently sourced from northern Nevada, appear in late Eocene fluvial deposits on the west slope of the Sierra Nevada, suggesting that any potential drainage divide along the arc had been breached (Cassel et al., in press). By ~30 Ma, exhumation was effectively complete (Fig. 4.12B), and the relict topography of the ancestral Sierra Nevada was subdued to near the baseline topography of the surrounding region. By this time, several westward-flowing paleovalleys carried

Oligocene ignimbrites from central Nevada across the western slopes of the Sierra

Nevada (Henry, 2008; Cassel et al., 2009). Ignimbrites of this age are preserved in paleovalleys up to 1.6 km deep scattered thoughout NW Nevada (Henry, 2008), but our results suggest that in the eastern part of the batholith, the interfluves of these valleys may have continued eroding as much as a few hundred meters (Fig. 4.12A) until basement rocks were thoroughly buried in the Miocene. (Note that subsequent Miocene burial of up to ~1 km limits the sensitivity of AFT modeling to this part of the thermal history.)

Potential Causes of Erosional Exhumation

Our results suggest that high relief and denudation persisted along the arc 197 axis for tens of millions of years after magmatism ceased, but they do not directly constrain whether the topographic feature responsible might have been a plateau margin or a freestanding drainage divide, either of which might produce localized erosional exhumation. For example, the high relief present at the edge of a plateau could result in similar exhumation patterns, as orographic precipitation and erosion is usually focused on the margins of the plateau, leaving little moisture for precipitation across and erosion of the interior of the plateau (Fig. 4.13A; e.g., Masek et al., 1994). In this case, the eastward shift in maximum exhumation rates would be related to the progressive topographic retreat of the plateau margin (Fig. 4.13A; cf. Masek et al., 1994). Although restoration of cross sections (e.g., Chase et al., 1998; DeCelles et al., 1995), and mineral thermobarometry (e.g., Camilleri et al., 1997; Lewis et al., 1999; McGrew et al., 2000) directly east of the study area supply evidence for thick (~50-65 km) crust in the hinterland of the Sevier thrust belt (Fig. 4.1), consistent with a high, isostatically supported plateau, this hypothesis is not considered applicable at this latitude for two reasons: (1) Reconstructed Late Cretaceous crustal thicknesses in NW Nevada are just

~40 km and only thin slightly (to ~34 km) westward (Colgan et al., 2006b; Lerch et al.,

2007). It therefore seems unlikely that the crust in this area could have supported the sort of high and abrupt plateau edge typically invoked in cases of focused plateau-margin erosion (e.g., Masek et al., 1994; Godard et al., 2009). (2) Along-strike correlations between intrusive age and the duration and tempo of exhumation, consistent with geologically-reasonable rates of lithospheric cooling, are not explained by a simple plateau-edge model. For these correlations to arise by coincidence would be particularly surprising given that this segment of the arc is curved and truncates the pre-Cretaceous 198

West East A erosion at plateau edge Nevadaplano

thick crust not actually observed here hot, weak crust

erosion at cooling arc

hot weak crust hot, buoyant crust alone unable to drive sufficient exhumation

C erosion at cooling arc Nevadaplano

hot, weak crust lower crustal flow?

Figure 4.13. Topographically exaggerated cartoons depict three scenarios that might be used to explain the observed pattern of erosion: (A.) Eastward-regressing exhumation could be interpreted as the response to focused erosion along the edge of a high plateau, but the crust does not appear to have been thick enough to support this hypothesis. (B.) The arc could be envisioned as an isolated drainage divide, supported by the bouy- ancy of still-hot arc lithosphere, but it is unclear how the necessary magnitude of exhumation could be generated in this case. (C.) Residual arc heat could leave a weak lower crust capable of accepting hot lower crustal material squeezed out from beneath an adjacent high plateau. 199 fabric of the continental margin, which generally exerts strong control on deformation in the region (Fig. 4.2).

The arguments above also support the hypothesis proposed here that the thermal structure of the magmatic arc played an important role in its exhumational history.

Exhumation could be related to high topography supported by the thermal buoyancy of the arc (Fig. 4.13B), but thermal expansion of a 50-60 km lithosphere can only account for a couple hundred meters of excess elevation at most (cf. Heuze, 1983). The total denudation resulting from this exponentially decaying excess elevation should be increased by isostatic rebound, but only by a factor of about ρm/(ρm - ρc)/e ≈ 2, where ρm and ρc are the density of the crust and mantle and e is the base of the natural logarithm

(Sleep, 1971). Compositional changes to more felsic, less dense material during magmatism might give a small added boost. Alternatively, the relict heat of the arc might be responsible for thermal weakening of the lithosphere, leading to enhanced crustal shortening during and after arc development, which in turn could lead to increased topographic relief and faster exhumation. Although tectonic deformation of this age has not been observed in supracrustal wall rocks of the batholith in NW Nevada (e.g.,

Compton, 1960; Wyld, 1996; Van Buer, submitted), ,material flow from the region shortening to the east into the base of the arc could lead to increased elevations with minimal associated deformation in the upper crust, similar to mechanisms proposed by

DeCelles and Coogan (2006) or by Royden et al. (1997) for eastern Tibet. In the case of the northernmost Sierra Nevada Batholith, we speculate that the inflow of lower crustal material could be modulated by the decaying geothermal gradient of the arc, explaining both the correlations between arc thermal structure and exhumation patterns and also the 200 amplification of the exhumation signal beyond what might be explained by thermal buoyancy alone (Fig. 4.13C). This model is also consistent with the evidence for a relatively high “Nevadaplano” to the east (e.g., Allmendinger, 1992; DeCelles, 2004;

Colgan and Henry, 2009) without actually requiring 50-65 km thick crust at the location of the arc (Fig. 4.13C). If the eroding topography of the arc was perched near the edge of a plateau to the east, drainage from between the batholith and the plateau would need to have either crossed the axis of the arc (for which there is no clear evidence until ~35 Ma;

Cassel et al., in press), or it could have flowed northwards towards the Hornbrook-

Ochoco basin (cf. Nilsen, 1984, 1993). Additional thermochronology data, particularly to the east of our study area, will be needed to constrain exhumation rates in the surrounding region before this or other models for exhumation can be tested fully.

SUMMARY

The exposure of the northernmost Sierra Nevada batholith afforded by normal faulting in the NW Basin and Range province provides access to samples from a broad range of Cretaceous to mid-Tertiary paleodepths in the batholith referenced to the regionally developed Tertiary unconformity, providing a consistent datum for interpreting these depths via analysis of maps and cross sections. Thermochronology data, including apatite fission track, apatite and zircon U-Th/He, biotite and K-feldspar 40Ar/39Ar, yield generally consistent thermal histories for each of the nine sample transects studied, across a region approximately 150 by 300 km, with apatite fission track and K-feldspar

40Ar/39Ar multi-diffusion domain modeling supplying the most robust constraints on temperature-time histories of samples. Differences in the thermal history of samples as a 201 function of depth in each transect are broadly interpretable in terms of a geothermal gradient that decreased gradually over time, reaching ~20–30 °C/km by the time Miocene extension began. With the exception of rapid cooling from high temperatures in some sections, which could be related to magmatic cooling, most of the measured thermal histories can be best interpreted in terms of the erosional denudation of the batholith.

Cooling related to erosional exhumation is demonstrated to last longer in the eastern part of the batholith, and the timing of exhumation is correlated with the intrusive age of basement rocks. The total duration of exhumation and the roughly exponentially decreasing rates of exhumation are consistent with the timescale of cooling for 50–60 km arc lithosphere. Used as a proxy for relative topographic relief, exhumation rates suggest that the relict topography of the magmatic arc waned rapidly after the cessation of magmatism at 90–85 Ma, but persisted until the mid-Tertiary (ca. 35 Ma), and the axis of greatest relief may have moved slowly eastwards with time. The progress of exhumation appears to be modulated by the slow cooling of the arc lithosphere, but the magnitude of exhumation is too great to be explained by thermal buoyancy forces alone. Crustal deformation and consequent uplift and/or addition of material at depth may have been localized along the arc because it was hotter and weaker than surrounding areas, possibly by allowing material from the east to be forced beneath the hot arc.

ACKNOWLEDGEMENTS

This research was funded by NSF Tectonics grant 0809226. Special thanks to

Matt Coble for help in the argon lab and help with argon data reduction, to Julie Fosdick for help with zircon dissolution for U-Th/He analysis, to Joe Colgan for use of his 202 samples, apatite fission track data, and data reduction scripts, and to Chris Gallagher for help with He extraction at UCSC. Special thanks to Elizabeth Miller for her review.

REFERENCES

Allmendinger, R.W., 1992, Fold and thrust tectonics of the western United States exclusive of the accreted terranes, in Burchfiel, B.C., Lipman, P.W., and Zoback, M.L., eds., The Cordilleran Orogen: Conterminous U.S.: The Geological Society of America, The Geology of North America, v. G-3, p. 583-607.

Armstrong, P.A., Ehlers, T.A., Chapman, D.S., Farley, K.A., and Kamp, P.J., 2003, Exhumation of the central Wasatch Mountains, Utah; 1, Patterns and timing of exhumation deduced from low-temperature thermochronology data: Journal of Geophysical Research, v. 108 (B3), 2172, doi:10.1029/2001JB001708.

Bingler, E.C., 1972, K-Ar dates from volcanic and plutonic rocks of the northern Wassuk Range, central Western Nevada: Isochron/West, v. 3, p. 31-32.

Bingler, E.C., 1978, Geologic map of the Schurz quadrangle: Reno, NV, Nevada Bureau of Mines and Geology, Map 60.

Brown, K., Stuck, R., and Hart, W.K., 2010, Geochronology and geochemistry of a Late Cretaceous granitoid suite, Santa Rosa Range, Nevada; Linking arc magmatism in northwestern Nevada to the Sierra Nevada Batholith: Abstract V23B-2445 presented at 2010 Fall Meeting, AGU, San Francisco, CA, 14 December.

Brueseke, M.E., and Hart, W.K., 2008, Geology and Petrology of the mid-Miocene Santa Rosa-Calico Volcanic Field: Bulletin – Nevada Bureau of Mines and Geology, Report 113, 83 pp.

203

Brueseke, M.E., and Hart, W.K., 2009, Intermediate composition magma production in an intracontinental setting; unusual andesites and dacites of the mid-Miocene Santa Rosa-Calico volcanic field, northern Nevada: Journal of Volcanology and Geothermal Research, v. 188, p. 197-213.

Busby, C.J., and Putirka, K., 2009, Miocene evolution of the western edge of the Nevadaplao in the central and northern Sierra Nevada—paleo canyons, magmatism, and structure: International Geology Review, v. 51, p. 670-701.

Cassel, E.J., and Graham, S.A., 2011, Paleovalley morphology and fluvial system evolution of Eocene- Oligocene sediments ('auriferous gravels'), northern Sierra Nevada, California: Implications for climate, tectonics, and topography, Geological Society of America Bulletin, v. 123, no. 9-10, p. 1699-1719.

Cassel, E.J., Grove, M., and Graham, S.A., in press, Eocene drainage evolution and erosion of the Sierra Nevada batholith across northern California and Nevada, American Journal of Science.

Cecil, M.R., Ducea, M.N., Reiners, P.W., Chase, C.G., 2006, Cenozoic exhumation of the northern Sierra Nevada, California, from (U-Th)/He thermochronolgy: Geological Society of America Bulletin, v. 118, p. 1481-1488.

Clark, M.C., and Royden, L.H., 2000, Topographic ooze: Building the eastern margin of Tibet by lower crustal flow: Geology, v. 28, p. 703-706.

204

Coble, M.A., Grove, M., and Calvert, A.T., 2011, Calibration of Nu-Instruments Noblesse multicollector mass spectrometers for argon isotopic measurements using a newly developed reference gas: Chemical Geology, v. 190, p. 75-87.

Coble, M.A., Scarberry, K., Grunder, A.L., and Wiegand, B.A., 2007, Chemical, isotopic, and Ar-geochronologic evidence for three-phase volcanic history of the northwest Basin & Range and High Lava Plains: Eos, Transactions of the American Geophysical Union, v. 88 (52), Fall Meeting Supplement, Abstract V51B-0573.

Coleman, D.S., and Glazner, A.F., 1997, The Sierra Crest magmatic event; rapid formation of juvenile crust during the Late Cretaceous in California: International Geology Review, v. 39, p. 768-787.

Colgan, J.P., 2005, Timing and Magnitude of Basin and Range Extension in Northwestern Nevada: Stanford University [Ph.D. thesis].

Colgan, J.P., Dumitru, T.A., and Miller, E.L., 2004, Diachroneity of Basin and Range extension and Yellowstone hotspot volcanism in northwestern Nevada: Geology v. 32, p. 121-124.

Colgan, J.P., Dumitru, T.A., McWilliams, M.O., and Miller, E.L., 2006a, Timing of Cenozoic volcanism and Basin and Range extension in northwestern Nevada; new constraints from the northern Pine Forest Range: Geological Society of America Bulletin, v. 118, p. 126-139.

Colgan, J.P., Dumitru, T.A., Reiners, P.W., Wooden, J.L., Miller, E.L., 2006b, Cenozoic tectonic evolution of the Basin and Range Province in northwestern Nevada: American Journal of Science, v. 306, p. 616-654.

Colgan, J.P., and Henry, C.D., 2009, Rapid middle Miocene collapse of the Mesozoic orogenic plateau in north-central Nevada: International Geology Review, v. 51, p. 920-961.

Colgan, J.P., Shuster, D.L., and Reiners, P.W., 2008, Two-phase Neogene extension in the northwestern Basin and Range recorded in a single thermochronology sample: Geology, v. 36, p. 631-634.

Colgan, J.P., Wyld, S.J., and Wright, J.E., 2010, Geologic map of the Vicksburg Canyon quadrangle, Humboldt County, Nevada: Nevada Bureau of Mines and Geology Map 169.

Compton, R.R., 1960, Contact metamorphism in Santa Rosa Range, Nevada: Geological Society of America Bulletin, v. 71, p. 1383-1416.

DeCelles, P.G., 2004, Late Jurassic to Eocene evolution of the Cordilleran thrust belt and foreland basin system, western U.S.A.: American Journal of Science, v. 304, p. 205

105-168.

DeCelles, P.G., and Coogan, J.C., 2006, Regional structure and kinematic history of the Sevier fold-and-thrust belt, central Utah: Geological Society of America Bulletin, v. 118, p. 841-864.

Dodson, M.H., 1973, Closure temperature in cooling geochronological and petrological systems: Contributions to Mineralogy and Petrology, v. 40, p. 259-274.

Ducea, M., 2001, The California arc: Thick granitic batholiths, eclogitic residues, lithospheric-scale thrusting, and magmatic flare-ups: GSA Today, v. 11, p. 4-10.

Dumitru, T.A., 1990, Subnormal Cenozoic geothermal gradients in the extinct Sierra Nevada magmatic arc: Consequences of Laramide and post-Laramide shallow- angle subduction: Journal of Geophysical Research, v. 95 (B4), p. 4925-4941.

Dumitru, T.A., 1993, A new computer-automated microscope stage system for fission- track analysis: Nuclear Tracks and Radiation Measurements, v. 21, p. 575-580.

Dumitru, T.A., Gans, P.B., Foster, D.A., and Miller, E.L., 1991, Refrigeration of the western Cordilleran lithosphere during Laramide shallow-angle subduction: Geology, v. 19, p. 1145-1148.

Egger, A.E., Colgan, J.P., and York, C., 2009, Provenance and paleogeographic implications of Eocene-Oligocene sedimentary rocks in the northwestern Basin and Range: International Geology Review, v. 51, p. 900-919.

Egger, A.E., Glen, J.M.G., and Ponce, D.A., 2010, The northwestern margin of the Basin and Range Province, part 2—Structural setting of a developing basin from seismic and potential field data: Tectonophysics, v. 488, p. 143-149.

Egger, A.E., and Miller, E.L., 2011, Evolution of the northwestern margin of the Basin and Range; The geology and extensional history of the Warner Range and environs, northeastern California: Geosphere, v. 7, p. 756-773.

Evernden, J.F., and Kistler, R.W., 1970, Chronology of emplacement of Mesozoic batholith complexes in California and western Nevada: U.S. Geological Survey Professional Paper 623.

Farley, K.A., Wolf, R.A., and Silver, L.T., 1996, The effects of long alpha-stopping distances on (U-Th)/He ages: Geochimica et Cosmochimica Acta, v. 60, p. 4223- 4229.

Faulds, J.E., Henry, C.D., Hinz, N.H., 2005, Kinematics of the northern Walker Lane; an incipient transform fault along the Pacific-North American Plate boundary: Geology, v. 33, p. 505-508. 206

Fitzgerald, P.G., Baldwin, S.L., Webb, L.E., and O’Sullivan, P.B., 2006, Interpretation of (U-Th)/He single grain ages from slowly cooled crustal terranes: A case study from the Transantarctic Mountains of southern Victoria Land: Chemical Geology, v. 225, p. 91–120, doi: 10.1016/j.chemgeo.2005.09.001.

Flowers, R.M., Ketcham, R.A., Shuster, D.L., and Farley, K.A., 2009, Apatite (U–Th)/He thermochronometry using a radiation damage accumulation and annealing model: Geochimica et Cosmochimica Acta, v. 73, p. 2347-2365.

Fosdick, J.C., and Colgan, J.P., 2008, Miocene extension in the East Range, Nevada; a two-stage history of normal-faulting in the northern Basin and Range: Geological Society of America Bulletin, v. 120, p. 1198-1213.

Galbraith, R.F., and Laslett, G.M., 1993, Statistical models for mixed fission-track ages: Nuclear Tracks and Radiation Measurements, v. 21, p. 459-470.

Gashawbeza, E.M., Klemperer, S.L., Wilson, C.K., and Miller, E.L., 2008, Nature of the crust beneath northwest Basin and Range Province from teleseismic receiver function data: Journal of Geophysical Research, v. 113 (B10), citation B10308.

Gastil, G., 1983, Mesozoic and Cenozoic granitic rocks of southern California and western Mexico: Memoir – Geological Society of America, v. 159, p. 265-275.

Gilbert, H., Jones, C., Owens, T.J., Zandt, G., 2007, Imaging Sierra Nevada lithospheric sinking: EOS, v. 88, p. 225 & 229.

Gleadow, A.J.W., Duddy, I.R., Green, P.F., and Lovering, J.F., 1986, Confined fission track lengths in apatite; a diagnostic tool for thermal history analysis: Contributions to Mineralogy and Petrology, v. 94, p. 405-415.

Godard, V., Cattin, R., and Lave, J., 2009, Erosional control on the dynamics of low- convergence rate continental plateau margins: Geophsical Journal International, v. 179, p. 763-777.

Green, P.F., Duddy, I.R., Laslett, G.M., Hegarty, K.A., Gleadow, A.J.W., and Lovering, J.F., 1989, Thermal annealing of fission tracks in apatite; 4. Quantitative modeling techniques and extension to geological time scales: Chemical Geology (Isotope Geoscience Section), v. 79, p. 155-182.

Grose, T.L.T., and Porro, C.T.R., 1989, Geologic Map of the Susanville 15-minute quadrangle, Lassen and Plumas Counties, California: Open-File Report – California Geological Survey Report 89-33.

Grove, M., Lovera, O., Harrison, M., 2003, Late Cretaceous cooling of the east-central Peninsular Ranges batholith (33°N): Relationship to La Posta pluton 207

emplacement, Laramide shallow subduction, and forearc sedimentation, in Johnson, S.E., Paterson, S.R., Fletcher, J.M., Girty, G.H., Kimbrough, D.L., and Martin-Barajas, A., eds., Tectonic evolution of northwestern Mexico and the southwestern USA: Boulder, Colorado, Geological Society of America Special Paper 374, p. 355-379.

Harrison, T.M., Heizler, M.T, Lovera, O.M., Chen W., and Grove, M., 1994, A chlorine disinfectant for excess argon released from K-feldspar during step-heating: Earth and Planetary Science Letters, v. 123, p. 95-104.

Heuze, F.E., 1983, High-temperature mechanical, physical, and thermal properties of granitic rocks—A review: International Journal of Rock Mechanics and Mining Sciences & Geomechanics Abstracts, v. 20, p. 3-10.

Hinz, N.H., Faulds, J.E., and Henry, C.H., 2009, Tertiary volcanic stratigraphy and paleotopography of the Diamond and Fort Sage Mountains; constraining slip along the Honey Lake fault zone in the northern Walker Lane, northeastern California and western Nevada: Special Paper – Geological Society of America, v. 447, p. 101-131.

Hourigan, J.K., Reiners, P.W., and Brandon, M.T., 2005, U-Th zonation-dependent alpha- ejection in (U-Th)/He chronometry: Geochimica et Cosmochimica Acta, v. 13, p. 3349- 3365.

House, M.A., Wernicke, B.P., Farley, K.A., and Dumitru, T.A., 1997, Cenozoic thermal evolution of the central Sierra Nevada, California, from (U-Th)/He thermochronometry: Earth and Planetary Science Letters, v. 151, p. 167-179.

House, M.A., Wernicke, B.P., and Farley, K.A., 1998, Dating topography of the Sierra Nevada, California, using apatite (U-Th)/He ages: Nature, v. 396, p. 66-69.

House, M.A., Wernicke, B.P., and Farley, K.A., 2001, Paleo-geomorphology of the Sierra Nevada, California, from (U-Th)/He ages in apatite: American Journal of Science, v. 301, p. 77-102.

Hudson, D.M., John, D.A., and Fleck, R.J., 2006, Geology, geochemistry, and geochronology of epithermal gold-silver deposits in the Seven Troughs District, Pershing County, Nevada, in Geology and Mineral Resources of the Trinity, Seven Troughs, and Kamma Ranges, West-Central Nevada: Geological Society of Nevada Spring 2006 Field Trip Guidebook, Special Publication 42, p. 110-126.

208

Hurford, A.J., and Green, P.F., 1983, The zeta age calibration in fission-track dating: Chemical Geology, v. 41, p. 285-317.

Johnson, M., 1977, Geology and mineral deposits of Pershing County, Nevada: Nevada Bureau of Mines and Geology Bulletin 89.

Jones, C.H., Farmer, G.L., Unruh, J, 2004, Tectonics of Pliocene removal of lithosphere of the Sierra Nevada, California, Geological Society of America Bulletin, v. 116 (11-12), p. 1408-1422.

Ketcham, R.A., Donelick, R.A., and Carlson, W.D., 1999, Variabiity of apatite fission- track annealing kinetics. 3. Extrapolation to geological time scales: American Mineralogist, v. 84, p. 1235-1255.

Ketcham, R.A., 2005, Forward and inverse modeling of low-temperature thermochronology data: Reviews in Mineralogy and Geochemistry, v. 58, p. 275- 314.

Klinger, M., 2005, Structral characteristics and U-Pb geochronological constraints of the Pine Nut Fault, Wassuk Range, west-central Nevada [M.S. thesis]: University of Idaho, Moscow, ID, 30 pp.

LeMasurier, W.E., 1965, Volcanic geology of Santa Rosa Range, Humboldt County, Nevada [Ph.D. thesis]: Stanford University, Stanford, CA, 126 pp.

Lerch, D.W., Klemperer, S.L., Stokoe, K.H., Menq, F.-Y., 2008a, Integration of the NEES T-Rex vibrator and PASSCAL Texan recorders for seismic profiling of shallow and deep crustal targets: Seismological Research Letters, v. 79, p. 41-46.

209

Lerch, D.W., Miller, E.L., McWilliams, M., and Colgan, J., 2008b, Tectonic and magmatic evolution of the northwestern Basin and Range and its transition to unextended volcanic plateaus; Black Rock Range, Nevada: Geological Society of America Bulletin, v. 120, p. 300-311.

Lerch, D.W., Klemperer, S.L., Egger, A.E., Colgan, J.P., Miller, E.L., 2010, The northwestern margin of the Basin-and-Range Province, part 1—Reflection profiling of the moderate-angle (~ 30 degree) Surprise Valley Fault: Tectonophysics, v. 488, p. 143-149.

Lide, D.R., ed., 1995, CRC Handbook of Chemistry and Physics, 76th edition. Boca Raton, Florida: CRC Press, Inc.

Lindgren, W., 1911, The Tertiary gravels of the Sierra Nevada of California, U. S. Geological Survey Professional Paper 73, 226 p.

Lovelock, E.C., 2010, Geology, Geochronology, and Paleobotany of the Moonlight and Susanville Fossil Flora Localities, Northern Sierra Nevada, California: [M.S. Thesis] Univeristy of California, Santa Barbara.

Lovera, O.M., Grove, M., Harrison, T.M., and Mahon, K.I., 1997, Systematic analysis of K-feldspar 40Ar/39Ar step-heating experiments I; Significance of activation energy determinations: Geochimica et Cosmochimica Acta, v. 61, p. 3171-3192.

Lovera, O.M., Grove, M., and Harrison, T.M., 2002, Systematic analysis of K-feldspar 40Ar/39Ar step-heating experiments II; Relevance of laboratory K-feldspar argon diffusion properties to nature: Geochimica et Cosmochimica Acta, v. 66, p. 1237- 1255.

Ludwig, K.R., 2001, Squid: A users’ manual: Berkeley Geochronology Center Special Publication 2, 19 p.

Ludwig, K.R., 2003, Isoplot 3.00, a geochronological tool-kit for Excel: Berkeley Geochronology Center Special Publication 4, 67 pp.

MacGinitie, H.D., 1941, A Middle Eocene flora from the central Sierra Nevada: Contributions to Palaeontology, Carnegie Institute of Washington Publication, v. 534, p. 1-169.

210

Martin, A.J., Wyld, S.J., Wright, J.E., and Bradford, J.H., 2010, The Lower Cretaceous King Lear Formation, northwest Nevada; Implications for Mesozoic orogenesis in the western U.S. Cordillera: Bulletin of the Geological Society of America, v. 122, p. 537-562.

Manganelli, S., 2009, Low-temperature argon diffusion in basaltic glass measured using image furnace methods: Stanford University [M.S. thesis]

Masek, J.G., Isacks, B.I., Gubbels, T.L., and Fielding, E.J., 1994, Erosion and tectonics at the margins of continental plateaus: Journal of Geophysical Research, v. 99 (B7), p. 13,941-13,956.

McDowell, F.W., McIntosh, W.C., and Farley, K.A., 2004, A precise 40Ar-39Ar reference age for the Durango apatite (U-Th)/He and fission-track dating standard: Chemical Geology, v. 214, p. 249-263.

McGrew, A.J., Peters, M.T., and Wright, J.E., 2000, Thermobarometric constraints on the tectonothermal evolution of the East Humboldt Range metamorphic core complex, Nevada: Geological Society of America Bulletin, v. 112, p. 45-60.

McKnight, B., 1984, Stratigraphy and sedimentology of the Payne Cliffs Formation, southwestern Oregon, in Oregon and California, Field Trip Guidebook—Pacific Section, vol. 42, Geology of the Upper Cretaceous Hornbrook Formation, edited by T. Nilsen, pp. 187 – 194, Soc. of Econ. Paleontol. and Mineral., Tulsa, Oklahoma.

Moores, E.M., Wakabayashi, J., and Unruh, J.R., 2002, Crustal-scale cross-section of the U.S. Cordillera, California and beyond, its tectonic significance, and speculations on the Andean Orogeny: International Geology Review, v. 44, p. 479-500.

Mulch, A., Graham, S.A., Chamberlain, C.P., 2006, Hydrogen isotopes in Eocene river gravels and paleoelevation of the Sierra Nevada: Science, v. 313, p. 87-89.

Nilsen, T.H., 1984, Stratigraphy, sedimentology, and tectonic framework of the Upper Cretaceous Hornbrook Formation, Oregon and California, in Nilsen, T.H., ed., Geology of the Upper Cretaceous Hornbrook Formation, Oregon and California: Pacific Section S.E.P.M., v. 42, p. 51-88.

Nilsen, T., 1993, Stratigraphy of the Cretaceous Hornbrook Formation, southern Oregon and northern California, U. S. Geol. Surv. Prof. Pap., P1521.

Noble, D.C., McKee, E.H., Smith, J.G., and Korringa, M.K., 1970, Stratigraphy and geochronology of Miocene volcanic rocks in northwestern Nevada: U.S. Geological Survey Professional Paper 700D, p. 23-32.

211

Oldow, J.S., 1984, Evolution of a late Mesozoic back-arc fold and thrust belt, northwestern Great Basin, U.S.A.: Tectonophysics, v. 102, p. 245-274.

Pierce, L.A., and Morgan, K.L., 1992, The track of the Yellowstone hotspot— Volcanism, faulting, and uplift, in Link, P.K., et al., eds., Regional geology of eastern Idaho and western Wyoming: Geological Society of America Memoir 179, p. 1-53.

Proffett, J.M., Jr., and Dilles, J.H., 1984, Geologic map of the Yerington District, Nevada: Nevada Bureau of Mines and Geology Map 77, scale 1:24 000, 1 sheet.

Ransome, F.L., 1909, Notes on some mining districts in Humboldt County, Nevada: U.S. Geological Survey Bulletin 414.

Reeg, H., 2008, Seismic structure of the crust and upper mantle of the Sierra Nevada, California: Univ. Colorado [M.S. thesis], 139 pp.

Reiners, P.W., Spell, T.L., Nicolescu, S., and Zanetti, K.A., 2004, Zircon (U-Th)/He thermochronometry: He diffusion and comparisons with 40Ar/39Ar dating: Geochimica et Cosmochimica Acta, v. 68, p. 1857-1887.

Renne, P.R., Swisher, C.C., Deino, A.L., Karner, D.B., Owens, T.L., DePaolo, D.J., 1998, Intercalibration of standards, absolute ages and uncertainties in 40Ar/39Ar dating: Chemical Geology, v. 145, p. 117-152.

Richter, F.M., Lovera, O.M., Harrison, T.M., and Copeland, P., 1991, Tibetan tectonics from 40Ar/39Ar analysis of a single K-feldspar sample: Earth and Planetary Science Letters, v. 105, p. 266-278.

Roberts, C.T., 1985, Cenozoic evolotion of the northwestern Honey Lake basin, Lassen County, California: Colorado School of Mines Quarterly, v. 80 (1), 64 pp.

Royden, L.H., Burchfiel, B.C., King, R.W., Wang, E., Zhiliang C., Feng S., Yuping L., 1997, Surface deformation and lower crustal flow in eastern Tibet: Science, v. 276, p. 788-790.

Rytuba, J.J., and McKee, E.H., 1984, Peralkaline ash flow tuffs and calderas of the McDermitt volcanic field, southeast Oregon and north central Nevada: Journal of Geophysical Research, v. 89, p. 8616-8628.

Saleeby, J.B., 1981, Ocean floor accretion and volcanoplutonic arc evolution of the Mesozoic Sierra Nevada, California, in Ernst, W.G., ed., Rubey Volume on the 212

Geotectonic Development of California: Englewood Cliffs, New Jersey, Prentice-Hall, p. 132-181.

Saleeby, J.B., and Busby-Spera, C.J., 1992, Early Mesozoic evolution of the western U.S. Cordillera, in Burchfiel, B.C., Lipman, P.W., and Zoback, M.L., The Cordilleran Orogen-- Conterminous United States: Decade of North American Geology, Geological Society of America, v. G-3, p. 107-168.

Saleeby, J., Saleeby, Z., Nadin, E., and Maheo, G., 2009, Step-over in the structure controlling the regional west-tilt of the Sierra Nevada microplate— eastern escarpment system to Kern Canyon system: International Geology Review, v. 51, p. 634-669.

Sleep, N.H., 1971, Thermal effects of the formation of Atlantic continental margins by continental break up: Geophysical Journal of the Royal Astronomical Society, v. 24, p. 325-350.

Smith, J.G., McKee, E.H., Tatlock, D.B., Marvin, R.F., 1971, Mesozoic granitic rocks in northwestern Nevada: A link between the Sierra Nevada and Idaho Batholiths: Geological Society of America Bulletin, v. 82, p. 2933-2944.

Stockli, D.F., Surpless, B.E., Dumitru, T.A., and Farley, K.A., 2002, Thermchronological constraints on the timing and magnitude of Miocene and Pliocene extension in the central Wassuk Range, western Nevada: Tectonics, v. 21 (4), 1028, doi: 10.1029/2001TC001295.

Surpless, B.E., 1999, Tectonic evolution of the northern Sierra Nevada-Basin and Range transition zone: A study of crustal evolution in extensional provinces [Ph.D. thesis]: Stanford University, 86 p.

213

Surpless, B.E., Stockli, D.F., Dumitru, T.A., Miller, E.L., 2002, Two-phase westward encroachment of Basin and Range extension into the northern Sierra Nevada: Tectonics, v. 21, doi. 10.1029/2000TC001257.

Unruh, J.R., Dumitru, T.A., and Sawyer, T.L., 2007, Coupling of early Tertiary extension in the Great Valley forearc basin with blueschist exhumation in the underlying Franciscan accretionary wedge at Mount Diablo, California: Geological Society of America Bulletin, v. 119, p. 1347-1367.

Van Buer, N.J., submitted, Preliminary Geologic Map of the Sahwave and Nightingale Ranges, Pershing County, Nevada. [version of Plates 1-8]

Van Buer, N.J., Miller, E.L., Dumitru, T.A., 2009, Early Tertiary paleogeologic map of the northern Sierra Nevada batholith and the northwestern Basin and Range: Geology, v. 37, p. 371-374.

Van Buer, N.J. and Miller, E.L., 2010, The Sahwave Batholith, NW Nevada: Cretaceous arc flare-up in a basinal terrane: Lithosphere, v. 2, p. 423-446.

Wagner, G.A., and Reimer, G.M., 1972, Fission track tectonics— The tectonic interpretation of fission track apatite ages: Earth and Planetary Science Letters, v. 14, p. 263-268.

Wakabayashi, J., and Sawyer, T.L., 2001, Stream incision, tectonics, uplift, and evolution of topography of the Sierra Nevada, California, Journal of Geology, v. 109 (5), p. 539-562.

Watanabe, K., Izawa, E., Kuroki, K., Honda, T., and Nakamura, H., 1991, Detection of confined 238U fission tracks in minerals and its application to geothermal geology: Annual Report of Tandem Accelerator Laboratory, Kyushu University, v. 3, p. 151-155.

Whitehill, C. S., 2009, Cenozoic evolution of the Shawave-Nightingale horst block, northwestern Basin and Range, Nevada, U.S.A.: Stanford University [Ph.D. 214

thesis].

Wolfe, J.A., 1994, Tertiary climatic changes at middle latitudes of western North America: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 108, p. 195- 205.

Wooden, J.L., R.E. Wells and J.A. Vance, 1999, Provenance of the Eocene Tyee Formation and related sandstones, western Oregon and Washington; detrital zircon evidence, Abstracts with Programs – Geological Society of America, v. 31 (7), p. 299.

Wyld, S.J., 1996, Early Jurassic deformation in the Pine Forest Range, northwest Nevada, and implications for Cordilleran tectonics: Tectonics, v. 15, p. 566–583.

Wyld SJ, Wright JE, 1997, Triassic-Jurassic tectonism and magmatism in the Mesozoic continental arc of Nevada; Classic relations and new developments, in Line, P.K., and Kowallis B.J., eds., Proterozoic to Recent Stratigraphy, Tectonics, and Volcanology, Utah, Nevada, Southern Idaho, and Central Mexico: Geological Society of America Field Trip Guide Book, Brigham Young University Geology Studies Vol. 42, Part 1, p. 197-224.

Yeend, W.E., 1974, Gold-bearing gravel of the ancestral Yuba River, Sierra Nevada, California: U.S. Geological Survey Professional Paper 722, 44 215

APPENDIX B: APATITE FISSION TRACK ANALYTICAL

PROCEDURES

All samples were analyzed at Stanford University. Apatite separates were generated using standard procedures, including rock crushing, grinding in a disc mill, washing, Frantz magnetic separation, and dense liquid separation techniques. Apatites were poured onto glass slides coated with a sugar/gum arabic mixture and mounted in epoxy bonded to a second slide, such that the sugar-coated slide could be removed, leaving the apatite grains near the top of a ~200 µm epoxy layer. The grains mounted in epoxy were ground halfway through using 2500 grit sandpaper and polished using 0.3

µm alumina suspension on a cloth lapidary wheel. Slides were cleaned ultrasonically with distilled water, and rinsed with isopropanol and acetone to remove any surface contamination before etching with 5.0 M nitric acid at 21°C for 20 seconds to reveal fission tracks. External mica detectors were affixed to the slides and pieces of CN5 dosimetry glass, which were stacked single-file in plastic tubes for irradiation. Three

CN5 glasses were included in each tube to monitor neutron flux. Samples were sent to the TRIGA reactor at Oregon State University and were irradiated by thermally activated neutrons. After irradiation, the mica detectors were etched in 48% hydrofluoric acid to reveal the induced fission tracks.

Fission tracks were counted on a Zeiss Axioskop optical microscope under transmitted or reflected light, as necessary, using a 100x air objective with 10x oculars and a 1.25x tube factor (total magnification 125,000x). External detector prints of the apatite grains were located using a Kinetek automated scanning stage (e.g., Dumitru

1993). Only grains oriented with the C-axis subparallel to the slide plane were counted 216 or measured for track lengths. Reported ages (Tables 4.2, B1; Fig. B1) use a zeta calibration factor of 337.3 ± 4.6 (cf. Hurford and Green, 1983). Fission-track central ages, after Galbraith and Laslett (1993), were calculated using an Excel program by J.

Colgan (e.g., Colgan et al., 2006b). Duplicate slides were irradiated with 252Cf fission fragments at the University of Melbourne, prior to etching in nitric acid, to generate a higher yield of confined fission tracks for measuring track length distributions (e.g.,

Watanabe et al., 1991). Track lengths were measured on confined tracks (both track-in- track and track-in cleavage) using a computer digitizing tablet and drawing tube calibrated with a stage micrometer (Dumitru, 1993). Both C-axis orientation and the size of the track apertures (Dpar and Dprp) were measured in grains where track lengths were measured, after the method of Ketcham et al. (1999).

Track length distributions (Fig. B2) were modeled using HeFTy software (Fig.

B3; Ketcham, 2005) using the annealing model and C-axis projection method of Ketcham et al. (2007ab). Dpar was used as the kinetic parameter and for estimating initial track length. A value of 0.893 was used as the ratio of length reduction in standard.

Constraints on the late Tertiary reheating history (Fig. B3) were constructed using the estimated unconformity paleodepths (Table 4.2) and the estimated thickness (or range of possible thickness) of the Tertiary strata (Table 4.1). The geothermal gradient during reheating was assumed to be in the range of ~22–37 °C/km, reaching 10 ± 5 °C at the surface, which encompasses previous estimates for the geothermal gradient in the western

Basin and Range at this time (Stockli et al., 2002; Colgan et al., 2006b). The duration of reheating is constrained to begin simultaneously with deposition of the oldest Tertiary strata, and end right after deposition of the youngest pre-extensional Tertiary strata 217

(Table 3.1; Fig. B3). All other parts of the thermal history are assumed to be monotonic cooling since the basement should have been cooling off due to erosional exhumation right up until the deposition of strata (Fig. B3). The high-temperature constraints in the models are generally only limited to be younger than the U-Pb or biotite age of the rock

(Fig. B3). Inverse modeling was run to evaluate at minimum 10,000 cooling paths, but additional cooling paths would be tried until at least 10 good paths (p > 0.5), or 100 acceptable paths (p > 0.05), were found (Fig. B3). TABLE B1. APATITE FISSION TRACK ANALYTICAL DATA

Sample Irradiation # xls. Spontaneaous Induced Dosimeter Age ± 1σ # lengths mean length ± 1σ * * * number number ρS NS ρI NI ρD ND (Ma) (μm)

7T1A SU079-1 29 0.235 476 3.441 6981 1.450 7106 16.5 ± 1.1 15 13.84 ± 0.74 7T3 SU079-2 30 1.022 1847 3.907 7061 1.466 7106 64.5 ± 2.4 122 12.57 ± 0.20 7T5 SU079-3 30 1.532 1888 5.637 6948 1.482 7106 67.6 ± 2.1 151 12.81 ± 0.20 BN01 SU079-4 30 0.574 830 2.142 3096 1.498 7106 67.4 ± 2.9 152 13.13 ± 0.13 BN03 SU079-5 33 1.131 827 4.792 3503 1.514 7106 58.5 ± 3.0 153 11.70 ± 0.23 BN05 SU079-6 13 1.066 575 5.203 2807 1.529 7106 51.6 ± 2.9 152 12.26 ± 0.19 BN07 SU079-7 27 0.771 1196 4.052 6287 1.545 7106 49.4 ± 1.8 156 11.22 ± 0.19 BN09 SU079-8 20 1.294 785 7.087 4299 1.561 7106 47.9 ± 2.1 152 12.16 ± 0.15 BN11 SU079-9 30 0.558 368 2.956 1948 1.577 7106 46.6 ± 3.6 151 12.09 ± 0.19 DD1 SU079-10 50 0.812 841 2.409 2494 1.593 7106 89.3 ± 5.2 0 n.d.† DD5 SU079-11 16 1.055 521 3.372 1665 1.609 7106 84.3 ± 4.5 150 13.75 ± 0.11 DD8 SU079-13 19 1.674 1378 6.154 5067 1.640 7106 74.8 ± 2.6 150 13.40 ± 0.11 NBR1 SU079-14 30 3.468 3833 12.175 13457 1.656 7106 79.1 ± 2.3 151 13.61 ± 0.10 NBR3 SU079-15 30 3.454 3200 12.540 11620 1.672 7106 77.2 ± 2.1 150 13.77 ± 0.08 PF01 SU079-16 30 1.742 1492 6.865 5879 1.688 7106 70.9 ± 3.4 150 13.17 ± 0.09 SE01 SU079-17 19 1.334 1251 5.778 5420 1.704 7106 66.3 ± 3.0 66 12.14 ± 0.29 SE04 SU079-18 30 1.784 1065 8.466 5053 1.719 7106 61.2 ± 2.7 158 12.42 ± 0.12 SE06 SU079-19 30 1.079 2102 5.731 11169 1.735 7106 55.1 ± 1.9 151 12.41 ± 0.16 SE08 SU079-20 14 1.230 998 6.760 5487 1.751 7106 53.5 ± 2.1 151 12.84 ± 0.12 SE10 SU079-21 30 0.648 1016 5.404 8468 1.767 7106 35.8 ± 1.6 80 12.51 ± 0.23 SE12 SU079-22 25 0.455 762 5.547 9286 1.783 7106 23.6 ± 1.3 113 12.73± 0.15 SL11 SU080-2 30 0.489 914 1.895 3539 1.450 6922 62.3 ± 3.0 154 13.36 ± 0.12 SL13 SU080-3 30 0.949 1700 3.935 7049 1.463 6922 59.8 ± 2.9 151 13.40 ± 0.09 SR15 SU080-6 16 0.071 75 1.077 1131 1.503 6922 17.5 ± 2.7 107 12.91 ± 0.19 SR20 SU080-8 30 0.326 557 1.208 2062 1.529 6922 68.9 ± 3.9 151 13.67 ± 0.12 SR22 SU080-9 44 0.258 370 1.117 1604 1.543 6922 59.7 ± 3.7 150 13.42 ± 0.09 SR24 SU080-10 30 0.331 478 1.419 2052 1.556 6922 60.9 ± 3.4 85 13.18 ± 0.17 SR26 SU080-11 32 0.319 427 1.716 2295 1.569 6922 49.9 ± 2.8 27 12.77 ± 0.33 WK01 SU080-13 9 0.814 57 1.743 122 1.596 6922 124 ± 20 0 n.d.†

*(cm-2 × 106) †Not determined. No Cf-irradiated slide. 218 219

7T1A-1 7T3 7T5 BN01-2 BN03-1 16.5 ± 1.1 Ma 64.5 ± 2.4 Ma 67.6 ± 2.1 Ma 67.4 ± 2.9 Ma 58.5 ± 3 Ma n = 29 n = 30 n = 30 n = 30 n = 33

BN05-1 BN07-1 BN09-1 BN11-2 DD1-2 51.6 ± 2.9 Ma 49.4 ± 1.8 Ma 47.9 ± 2.1 Ma 46.6 ± 3.6 Ma 89.3 ± 5.2 Ma n = 13 n = 27 n = 20 n = 30 n = 50

DD5-1 DD8-1 NBR1-4 NBR3-2 PF01-1 84.3 ± 4.5 Ma 74.8 ± 2.6 Ma 79.1 ± 2.3 Ma 77.2 ± 2.1 Ma 70.9 ± 3.4 Ma n = 16 n = 19 n = 30 n = 30 n = 30

SE01-1 SE04-3 SE06-2 SE08-2 SE10-1 66.3 ± 3 Ma 61.2 ± 2.7 Ma 55.1 ± 1.9 Ma 53.5 ± 2.1 Ma 35.8 ± 1.6 Ma n = 19 n = 30 n = 30 n = 14 n = 30

SE12-1 SL-11 SL-13 SR15-2 SR-20 23.6 ± 1.3 Ma 62.3 ± 3 Ma 59.8 ± 2.9 Ma 17.5 ± 2.7 Ma 68.9 ± 3.9 Ma n = 25 n = 30 n = 30 n = 16 n = 30

Sample number SR-22 SR24-4 SR26-2 WK01-1 200 59.7 ± 3.7 Ma 60.9 ± 3.4 Ma 49.9 ± 2.8 Ma Age (Ma ± 1σ) 124.5 ± 20.1 Ma 150 n = 44 n = 30 n = 32 n = 9 Number of grains 100 Age (Ma) 70 2σ error 50 40 30 All plots at same scale. 20 1σ uncertainty 10 20 5 3 2 5 0 20 40 60 80 Precision Index Figure B1. Radial plots (Galbraith, 1990) showing apatite fission-track grain ages. 220

7T1A-1 7T3 7T5 BN01-2 BN03-1 16.5 ± 1.1 64.5 ± 2.4 67.6 ± 2.1 67.4 ± 2.9 58.5 ± 3 13.84 ± 0.74 12.57 ± 0.2 12.8 ± 0.2 13.13 ± 0.13 11.7 ± 0.23 n = 15 n = 122 n = 151 n = 152 n = 153

BN05-1 BN07-1 BN09-1 BN11-2 DD5-1 51.6 ± 2.9 49.4 ± 1.8 47.9 ± 2.1 46.6 ± 3.6 84.3 ± 4.5 12.26 ± 0.19 11.22 ± 0.19 12.16 ± 0.15 12.09 ± 0.19 13.75 ± 0.11 n = 152 n = 156 n = 152 n = 151 n = 150

DD8-1 NBR1-4 NBR3-2 PF01-1 SE01-1 74.8 ± 2.6 79.1 ± 2.3 77.2 ± 2.1 70.9 ± 3.4 66.3 ± 3 13.4 ± 0.11 13.61 ± 0.1 13.77 ± 0.08 13.17 ± 0.09 12.14 ± 0.29 n = 150 n = 151 n = 150 n = 150 n = 66

SE04-3 SE06-2 SE08-2 SE10-1 SE12-1 55.1 ± 1.9 53.5 ± 2.1 35.8 ± 1.6 23.6 ± 1.3 61.2 ± 2.7 12.41 ± 0.16 12.84 ± 0.12 12.51 ± 0.23 12.73 ± 0.15 12.42 ± 0.12 n = 151 n = 151 n = 80 n = 113 n = 158

SL-11 SL-13 SR15-2 SR-20 SR-22 62.3 ± 3.0 59.8 ± 2.9 17.5 ± 2.7 68.9 ± 3.9 59.7 ± 3.7 13.36 ± 0.12 13.40 ± 0.09 12.91 ± 0.19 13.67 ± 0.12 13.42 ± 0.09 n = 154 n = 151 n = 106 n = 151 n = 150

SR24-4 40 SR26-2 Sample number 60.9 ± 3.4 49.9 ± 2.8 13.18 ± 0.17 12.77 ± 0.33 Fission-track age (Ma ± 1σ) n = 85 n = 27 Mean track length (μm ± s.e.) 20 Number of tracks measured

Proportion of tracks (%) 0 All plots at same scale. 0 5 10 15 20 Track length (μm) Figure B2. Apatite fission-track length distributions. 221 Time before present (millions of years) 120 100 80 60 40 20 100 80 60 40 20 0 0 PF-01 SR-15 50 best-fit model 100 good results 150 acceptable results model constraints JC01-PF14 BN-01 50

100

150

JC02-PF44 BN-03 50

100

150

JC01-PF13 JC02-BR2 50

100

150

JC00-SR4 BN-05 50

100

150

SR-20 BN-07 50

100 Temperature (°C) Temperature 150

SR-22 BN-09 50

100

150

JC00-SR9 50 JC02-BR3

100

150

SR-24 50 BN-11

100

150

SR-26 7T-1A 50

100

150

200 Figure B3. Apatite fission track length modeling constraints and results. 222

Time before present (millions of years) 120 100 80 60 40 20 100 80 60 40 20 0 0 7T-3 SL-11 50

100

150

7T-5 SL-13 50

100

150

NBR-3 SE-01 50

100

150

JC02-BL5 SE-04 50

100

150

JC02-BL6 SE-06 50

100

150

JC02-BL7 SE-08 50

100 Temperature (°C) Temperature 150

NBR-1 SE-10 50

100

150

DD-5 SE-12 50

100

150

DD-8 50

100

150

200 Figure B3. (Continued) 223

APPENDIX C: APATITE AND ZIRCON U-Th/He ANALYTICAL

PROCEDURES

Apatite and zircon crystals separated by standard mineral separation procedures were hand-selected for large, euhedral, inclusion-free (or nearly so) grains that were not cracked or broken. Micrometer-calibrated optical microphotographs of each grain (down both a-axes for zircon crystals) were taken at 300x so that grain dimensions could be measured with image-editing software (Tables C1, C2). Each grain was packed in an

~0.8 mm square niobium packet. Helium was extracted with a diode laser and measured by 3He isotope dilution on a Pfeiffer Prisma quadrupole mass spectrometer at the

University of Santa Cruz (UCSC). Grains (still in packet) were spiked with 229Th and

233 U. Apatite grains (in Nb) were dissolved and diluted in HCl-HNO3 solution at UCSC.

Zircon grains (in Nb) were dissolved at Stanford in HF-HNO3 solution in Teflon microvials within Teflon-lined Parr bombs at 225°C for 72 hours and diluted for ICP use directly (no HF evaporation procedure). Uranium and thorium isotopes, as well as a survey scan of other elements, were measured on a Thermo XSERIES 2 quadrupole ICP-

MS at UCSC (Tables C1, C2). Samarium concentrations were below detection levels.

Alpha-ejection corrections were applied modeling apatite crystals as cylinders (Farley et al., 1996) and zircon crystals as bipyramidal terminated tetragonal prisms (Hourigan et al., 2005). Ages of Durango apatite and Fish Canyon Tuff zircon standards overlap with published reference ages (McDowell et al., 2004; Renne et al., 1998) within analytical precision (Tables C1, C2). In general, however, ages of the unknown samples do not overlap within analytical precision of each other—the standard deviation of each population is reported as the uncertainty in Table 4.2. 224

Apatite U-Th/He data were modeled with HeFty software (Ketcham, 2005) using the parameters and methods of the radiation damage accumulation and annealing model

(RDAAM) of Flowers et al. (2009). For modeling purposes, apatite was modeled as a

2 sphere with D0 = 0.6071 cm /s, containing traps of 34 kJ/mol energy, with the parameters

Ψ = 1×10-13 and Ω = 1 × 10-22, annealing according to the fission-track annealing model of Ketcham et al. (2007b), with rmr0 = 0.83 (Flowers et al., 2009). It was not possible to fit all apatite U-Th/He and fission track data for each sample simultaneously, and only in one case (sample BN01; Fig. 4.7C) was it possible to generate an inverse model that even satisfied all the apatite U-Th/He ages for a sample using the default value for activation energy. These scattered results could simply stem from analytical problems with this relatively new dating technique. Alternatively, U-Th zoning (e.g., Fitzgerald et al., 2006) or differences in the activation energy among the grains (e.g., Unruh et al.,

2007) may account for the differences in age (note that grain size and radiation damage effects are already accounted for in the RDAAM; Flowers et al., 2009). Visual observation of grains during fission-track analysis revealed that a majority of grains were significantly zoned (~ 56% among all apatite samples used for U-Th/He dating), almost always with narrow (~10 µm) high-U rims around lower-U cores, which would be expected to increase alpha ejection as well as diffusive He loss and result in anomalously young ages (e.g., Fitzgerald et al., 2006). Some of the relatively youngest grains (0062-

SE04, 0072-DD1 and 0076-BN05; Table C1) were from samples in which some grains were particularly strongly zoned, and they had significantly higher concentrations of U and Th than other grains, consistent with the presence of highly radiogenic rims. Zoning 225 may account for some of the scatter among all of the U-Th/He ages, and suggests that overly young ages may be particularly suspect.

Additionally, small differences in activation energy may be magnified into significant age differences by very slow cooling through the apatite He partial retention zone (as indicated by fission-track–length models); Colgan et al. (2006b) and Unruh et al. (2007) suggested a similar mechanism for generating a spread of ages in cases of reheating to temperatures in the He partial retention zone. To test this hypothesis, U-

Th/He results were forward-modeled separately from the fission-track data in HeFty to test whether the spread in ages might be plausibly generated with minor changes in activation energy (cf. Unruh et al., 2007). Forward models were based on the fission- track inverse modeling, but were manually adjusted for the best fit where the cooling paths cross the temperatures associated with the He partial retention zone. The modeled undamaged-lattice activation energy (EL of Flowers et al., 2009) for each grain was then changed until its modeled and measured ages matched. With the exception of the three anomalously young grains mentioned above, the range of EL needed to fit the forwards models is 118.6–132.4 kJ/mol (standard deviation: 3.4 kJ/mol; Fig. C1). Flowers et al.

(2009) constrain mean EL to be in the range of 119–122.3 kJ/mol (Fig. C1). However, the full natural range of EL should be approximated by the range in experimentally- derived activation energy (Ea) for samples with low effective fission track density (eρs),

5 2 which spans ~118–131 kJ/mol for eρs < 2 × 10 tracks/cm (Flowers et al., 2009), similar to the range modeled in our samples (Fig. C1).

Zircon U-Th/He closure temperatures shown in Fig. 4.7 were calculated by the formula of Dodson (1973), using the diffusion parameters reported by Reiners et al. 226

(2004) and cooling rates estimated from the best-fit thermal histories shown in Fig. 4.7.

The natural variation of zircon He diffusion parameters (Ea = 169 ± 3.8 kJ/mol; D0 =

+0.87 2 0.46 -0.30 cm /s; Reiners et al., 2004) suggests that 2σ errors on the closure temperatures calculated for zircon U-Th/He ages are on the order of ± 40°C if sample diffusion parameters are not measured. Such large uncertainties in the diffusion parameters make detailed diffusion modeling unproductive in this case, but could account for some of the spread in our zircon U-Th/He ages (Table C1). However, a surprising number of zircon U-Th/He single grain ages were younger than apatite-fission-track ages from the same sample, which is not consistent with any choice of diffusion parameters

(Table C1). These ages were not used. U-Th zonation and analytical problems (e.g., with our zircon dissolution procedure) could contribute to these age errors. 227

TABLE C1. ZIRCON U-Th/He ANALYTICAL DATA

* Grain number U Th He mass width 1 width 2 length tip FT (ppm) (ppm) (nmol/g) (µg) (µm) (µm) (µm) (µm) age, ± 2σ (Ma)

0001-FCT 470 228 56.4 3.8 69 67 225 36 0.73 27.3 ± 1.0† 0002-FCT 424 179 57.5 11.1 108 94 304 52 0.81 28.0 ± 0.8† 0003-FCT 462 230 50.7 3.3 67 66 203 32 0.72 25.2 ± 1.0† 0011-7T5 389 172 1.71 5.9 84 65 278 32 0.75 0.98 ± .04§ 0012-7T5 330 178 105 5.0 93 94 185 46 0.77 66.8 ± 2.4# 0013-7T5 139 60.0 45.1 5.2 103 89 177 42 0.78 69.3 ± 2.5 0014-7T5 181 108 63.5 3.9 125 75 149 44 0.76 74.4 ± 3.0 0015-BN01 118 36.3 37.0 30.9 178 182 325 90 0.88 61.0 ± 2.4# 0016-BN01 281 103 74.1 7.5 108 74 251 36 0.78 56.9 ± 2.0# 0017-BN01 0.1 1.5 0.13 7.0 96 93 240 53 0.77 63.6 ± 7.4** 0019-BN11 175 69.2 53.2 16.0 123 129 305 67 0.84 60.8 ± 2.3 0020-BN11 125 51.5 32.7 4.2 115 63 205 62 0.75 58.8 ± 2.5 0021-BN11 6.8 3.9 29.7 4.7 101 51 250 40 0.73 900 ± 39** 0022-BN11 110 44.7 30.1 7.4 103 87 235 42 0.79 57.8 ± 2.6 0023-PF01 177 60.5 64.7 12.3 126 109 257 48 0.83 75.2 ± 3.4 0024-PF01 215 60.6 94.2 7.3 87 75 290 36 0.77 97.3 ± 3.7 0025-PF01 212 83.9 57.9 2.5 72 60 163 29 0.71 64.9 ± 2.5# 0026-PF01 268 70.7 78.7 7.4 110 82 234 43 0.79 64.1 ± 2.2# 0027-JCPF1 81.5 49.4 17.1 70.8 175 173 606 76 0.89 38.0 ± 3.4 0028-JCPF1 149 50.4 18.9 18.9 147 119 313 59 0.85 25.5 ± 1.2 0029-JCPF1 113 45.2 11.7 7.8 108 97 225 49 0.80 21.8 ± 0.8 0030-JCPF1 133 46.9 25.6 16.7 124 121 315 57 0.84 38.8 ± 1.3 0031-SE12 475 125 117 13.5 105 103 275 48 0.84 50.8 ± 1.9 0032-SE12 529 124 116 14.1 113 113 298 45 0.83 46.3 ± 1.8 0033-SE12 368 139 134 9.2 100 99 256 41 0.80 76.3 ± 2.9 0034-SE12 1070 234 257 12.5 134 126 214 40 0.83 50.4 ± 1.8 0035-SE01 1580 1266 509 12.5 102 96 330 43 0.81 61.5 ± 2.1# 0036-SE01 835 439 259 12.2 106 107 288 41 0.82 61.9 ± 2.1# 0037-SE01 1120 373 405 4.6 103 92 161 43 0.77 79.1 ± 2.7 0038-SE01 453 141 140 7.6 108 88 234 47 0.79 66.6 ± 2.2 0039-SR15 374 134 101 5.2 83 87 224 53 0.77 59.6 ± 2.1 0040-SR15 327 121 91.9 5.5 99 86 203 48 0.78 61.1 ± 2.3 0041-SR15 402 142 105 6.02 82 89 252 57 0.77 57.3 ± 2.1 0042-SR15 305 121 62.6 5.71 97 88 197 41 0.78 44.3 ± 1.9 0043-SR20 1.4 0.1 266 6.71 74 78 315 47 0.76 12200 ± 1600** 0044-SR20 663 194 136 3.02 56 73 205 35 0.71 49.7 ± 2.1# 0045-SR20 1860 306 315 2.79 63 59 182 15 0.70 42.7 ± 1.8# 0046-SR20 521 199 164 3.36 85 69 193 52 0.74 72.1 ± 2.5 0047-WK13 617 224 135 7.89 92 101 228 33 0.79 46.7 ± 1.5 0048-WK13 451 132 104 4.20 82 79 201 46 0.75 52.4 ± 1.8 0049-WK13 681 232 135 3.25 83 66 176 36 0.73 45.9 ± 1.5 0050-WK13 410 161 102 6.06 106 94 188 42 0.79 53.1 ± 1.9

* FT alpha-ejection correction factor of Farley et al. (1996). †cf. Fish Canyon Tuff 40Ar/39Ar reference age of 28.02 ± 0.28 Ma (Renne et al., 1998) §He extraction problems. Result not used. #Younger than apatite fission track age from the same sample. Result not used. **Very low U, Th indicates problem (material lost?). Result not used. 228

TABLE C2. APATITE U-Th/He ANALYTICAL DATA

* Grain number U Th He mass radius length F T age, ± 2σ (ppm) (ppm) (nmol/g) (µg) (µm) (µm) (Ma)

0051-BN01 16.6 19.5 3.49 9.8 70 238 0.79 38.0 ± 2.0 0052-BN01 14.7 22.3 3.23 5.5 57 204 0.75 39.7 ± 1.8 0053-BN01 18.7 26.8 3.94 4.3 45 262 0.71 40.8 ± 1.7 0054-BN01 22.2 32.5 4.73 2.9 45 173 0.69 41.9 ± 2.4 0055-DD8 21.0 35.1 6.02 11.2 66 306 0.79 47.8 ± 1.9 0056-DD8 33.5 55.7 12.1 7.8 66 217 0.78 61.1 ± 2.1 0057-DD8 35.7 48.8 8.82 3.8 49 195 0.72 47.9 ± 2.0 0058-DD8 27.5 39.3 7.88 4.9 55 197 0.74 53.1 ± 2.3 0059-SE04 47.6 71.3 7.17 2.0 40 154 0.66 31.1 ± 1.4 0060-SE04 32.1 40.7 7.25 2.1 42 143 0.67 48.0 ± 2.0 0061-SE04 0.0 0.6 0.00 0.9 30 119 0.43 0† 0062-SE04 54.9 63.0 4.10 0.7 28 105 0.53 20.4 ± 1.3§ 0063-SL11 13.6 20.2 3.86 8.3 59 290 0.77 50.4 ± 1.9 0064-SL11 12.4 15.9 2.40 1.8 35 175 0.63 43.3 ± 3.5 0065-SL11 16.5 22.1 3.29 2.9 42 196 0.69 40.6 ± 1.7 0066-SL11 12.9 20.8 4.46 3.9 46 225 0.71 64.9 ± 2.9 0067-SR22 8.0 12.2 2.02 3.7 44 232 0.70 48.9 ± 2.8 0068-SR22 7.9 12.7 1.81 2.6 41 195 0.67 45.2 ± 3.1 0069-SR22 8.8 11.3 1.97 5.1 47 277 0.72 43.7 ± 1.6 0070-SR22 11.1 15.4 3.27 3.7 42 249 0.69 58.8 ± 2.7 0071-DD1 20.9 31.2 7.40 5.1 56 197 0.74 64.6 ± 2.3 0072-DD1 67.2 93.0 8.46 2.5 44 154 0.68 25.6 ± 1.2§ 0073-DD1 22.7 34.1 7.00 3.7 48 195 0.71 58.5 ± 2.3 0074-DD1 10.2 11.9 2.63 6.3 54 261 0.75 49.3 ± 2.2 0075-BN05 24.5 32.3 5.39 13.5 78 268 0.81 38.0 ± 2.1 0076-BN05 136 79.7 19.0 1.4 35 137 0.62 36.3 ± 1.3§ 0077-BN05 10.4 20.0 2.91 2.8 47 151 0.69 51.0 ± 2.6 0078-BN05 33.5 44.0 7.24 4.0 50 191 0.72 41.9 ± 2.0 0079-SR20 10.6 13.4 3.23 9.9 64 288 0.79 54.9 ± 2.4 0080-SR20 7.4 10.3 1.91 4.7 56 179 0.74 48.1 ± 2.0 0081-SR20 6.7 8.6 1.76 17.3 92 246 0.83 44.8 ± 1.4 0082-SR20 7.0 7.5 2.36 28.6 87 452 0.84 58.8 ± 1.7 0083-7T3 25.6 67.5 7.09 4.3 53 188 0.73 43.2 ± 1.3 0084-7T3 24.5 79.8 7.44 4.0 53 173 0.72 43.7 ± 1.6 0085-7T3 25.4 66.7 6.18 3.1 41 225 0.68 40.8 ± 1.6 0086-7T3 27.0 67.6 6.47 1.6 37 145 0.63 44.1 ± 2.2 0087-7T5 35.6 74.9 7.62 2.1 38 175 0.65 40.6 ± 1.4 0088-7T5 27.4 65.4 8.31 8.1 62 255 0.77 46.2 ± 1.8 0089-7T5 27.8 69.6 7.53 8.1 64 236 0.77 40.5 ± 1.4 0090-7T5 25.2 64.0 6.40 3.6 47 197 0.70 41.5 ± 1.4 0091-DUR n.d.# n.d.# n.d.# n.d.# n.d.# n.d.# 1.00 32.9 ± 1.5** 0092-DUR n.d.# n.d.# n.d.# n.d.# n.d.# n.d.# 1.00 32.1 ± 1.1** 0093-DUR n.d.# n.d.# n.d.# n.d.# n.d.# n.d.# 1.00 31.8 ± 1.3** 0094-DUR n.d.# n.d.# n.d.# n.d.# n.d.# n.d.# 1.00 32.0 ± 1.5** 0095-DUR n.d.# n.d.# n.d.# n.d.# n.d.# n.d.# 1.00 32.2 ± 1.0**

* FT alpha-ejection correction factor of Farley et al. (1996). †Grain probably lost from packet before analysis. Result not used. §High U-Th and low age; probably has narrow, high U-Th rim. Result not used. #Not determined. Size not measured because alpha-ejection assumed minimal for Durango standards. **cf. Durango standard 40Ar/39Ar reference age of 31.44 ± 0.18 Ma (McDowell et al., 2004) 229

80 A.

40 Apparent age (Ma)

B. 130 (kJ/mol) L 125 natural range of EL (?)

120 preferred range of EL

115 Fitted E 7T-3 7T-5 SR-20 SR-22 BN-05 DD-1 DD-8 SL-11 SE-04 Figure C1. (A.) Apatite (U-Th)/He ages for samples, labeled at bottom, that did not give acceptable inverse model results, showing 2σ errors. Grain ages from the same samples frequently do not overlap within error. (B.) Consistent forward models (Fig. 4.7) can be fit using the relatively narrow range of undamaged-lattice activation energies (EL) shown. These results exceed the bounds for mean EL suggested by Flowers et al. (2009), shown as a dark grey band, but approximate the apparent natural range of EL (see text), shown in light grey. 230

APPENDIX D: BIOTITE AND K-FELDSPAR 39Ar/40Ar ANALYTICAL

PROCEDURES

Biotite and K-feldspar were picked directly from small, hammer-crushed rock samples. Biotite was selected for euhedral, unaltered-looking, unfolded, and with no visible inclusions, about 20 per sample (~2 mg total) using a reflected light microscope.

Microcline was hand-picked under clove oil (index of refraction between microcline and plagioclase or quartz) using a petrographic microscope. Grains on the order of 0.5 mm in dimension (1-10 grains per sample, total mass ~2–4 mg), with as few inclusions as possible, were selected to provide enough argon for detailed stepheating work, and additionally to ensure diffusion information was not lost from the largest diffusion domains. In some samples, however, particularly in the tonalitic samples BN-11 and 7T-

5, K-feldspar takes a poikilitic habit, and it was not possible to find crystals of sufficient size that did not contain some amount of intergrown quartz or plagioclase inclusions.

Biotite and microcline grains were packed in ~5 mm copper foil packets within evacuated quartz-glass tubes alongside Taylor Creek sanidine flux monitors (28.34 ± 0.09 Ma;

Renne et al., 1998) as well as KCl, K2SO4, and CaF2 calibration salts (cf. Coble et al.,

2011). The samples were loaded into unshielded aluminum irradiation vessels and irradiated at McMaster Reactor for approximately 10 hours.

After irradiation, biotite and monitor grains were unpacked and loaded into a copper sample tray. Each sample was divided between 6–9 wells holding a few grains each, with each row in Table D1 representing fusion of the grains in a single well using a

CO2 laser. For the step-heating experiments, each copper packet containing irradiated K- feldspar grains was enclosed in an 8 mm metal foil packet along with a K-type 231 thermocouple. For samples WK-15 and 7T-5, these outer packets were made of copper; for sample PF-11, titanium; and for samples SE-12, SR-17, and WK-15, molybdenum.

Each packet was suspended by the thermocouple wires in vacuo behind a 23 mm pyrex or sapphire viewport. A 300 W projector lamp with gold-coated reflector, regulated by a

PID controller attached to the thermocouple, was mounted just outside the viewport for heating the sample, using an arrangement very similar to that of Manganelli (2009).

Stable temperatures (within ± 1°C) up to 900°C could typically be reached within 30 s of adjusting the target temperature (cf. Manganelli, 2009). Note that the highest temperatures could be achieved only using the Ti or Mo packets, which absorb much more strongly in the infrared than Cu (e.g., Lide, 1995). The use of titanium for sample

PF-11 resulted in a high background of atmospheric argon (Table D2), which had presumably been trapped within the titanium. The highest temperatures achieved were strictly beyond the design capability of the viewports used, resulting in a high risk of vacuum failure. At the conclusion of the stepheating experiments, the K-feldspar grains were unpacked and loaded into wells of a copper sample tray for laser fusion.

After each laser fusion or heating step, sample gas was gettered for 8 minutes before admission to the mass spectrometer. Isotopic ratios of 40Ar, 39Ar, 38Ar, 37Ar, and

36Ar were measured using a Nu Instruments Noblesse multicollector mass spectrometer in static mode, using approximately the same extraction setup used in Coble et al. (2011;

Tables D1, D2). 40Ar, and occasionally 39Ar if large signals were expected, were measured using a Faraday cup detector (1011 ohm resistor); otherwise measurements were made using two discrete-dynode electron multipliers in ion-counting mode on the low- mass side of the detector array (cf. Coble et al., 2011). Mass fractionation and detector 232 bias were corrected by daily (or more frequent) measurements on air, a reference gas mixture containing subequal proportions of 40Ar, 38Ar, and 39Ar (4.766 : 0.751 : 1), and mixtures thereof, using the method of Coble et al. (2011).

Measurement of the Taylor Creek sanidine flux monitors revealed statistically significant flux gradients along the length of the irradiation tubes (up to 6‰), so the J- factors calculated for the unknowns were interpolated via linear regression among the flux monitors in each tube. The following isotopic correction factors were used:

40 39 38 39 36 37 Ar/ ArK = 0.0288 ± 0.0006, Ar/ ArK = 0.0128 ± 0.0003, Ar/ ArCa = 0.00026 ±

39 37 38 39 0.00004, Ar/ ArCa = 0.00066 ± 0.0002, and ArCl/ ArK = 2.77 ± 0.30. Low temperature steps of the K-feldspar step-heating data were chlorine-corrected using the method of Harrison et al. (1994), by calculating and substracting the Cl-correlated 40Ar content estimated from isothermal heating steps, which is assumed to be derived from fluid-inclusion-hosted excess argon (Table D2). Cl/K for each step is estimated from

38 39 38 Ar/ ArK after correcting for atmospheric and K-produced Ar. For some samples, the excess-argon–chlorine ratio could not be very precisely determined; however, in these cases, the recognition of large uncertainty for the early, Cl-rich steps (Table D2) still led to improved modeling of the data because these imprecise steps became less heavily weighted.

K-feldspar step-heating data were modeled using the method and software of

Lovera et al. (1997, 2002) to recover both diffusion parameters and thermal history information from each sample, assuming the presence of multiple, non-interacting diffusion domains (Figs. 4.7, D1). Heating steps below 400°C yielded little 39Ar and were omitted from the thermal modeling (Table D2). Problems encountered during 233 modeling of the data (Table D2) include irregularities in the diffusion data (samples PF-

11, SE-12, and WK-15) and non-monotonic age spectra (samples SE-12, SR-17, and the final heating steps of WK-15). Problems with the diffusion data could generally be traced to loss of unmeasured argon, either because of malfunctioning or misoperated valves, e.g., the first 533°C step for sample WK-15 (Table D2, Figure D1); or, in the case of sample PF-11, because the heating schedule was interrupted at 600°C for additional

250°C baking due to high 36Ar counts from the Ti packet. Non-monotonic age spectra imply that some of the assumptions of multi-diffusion-domain modeling may not be met by that sample; for example, large domains may have recrystallized into smaller domains during low-temperature adularia alteration (e.g., Lovera et al., 2002). Forward modeling suggests that meaningful thermal history information may still be recovered in these cases for temperatures below that of the recrystallization event (Lovera et al., 2002).

Therefore, for samples with non-monotonic age spectra, only the initial, age-monotonic gas release was modeled (higher temperature steps and the laser fusion gas release were averaged together, weighted by 39Ar). Sample PF-11 yielded a particularly irregular age spectrum, which may be partly the result of applying major atmospheric corrections to heating steps that were as little as 1.87% radiogenic 40Ar at the highest temperatures

(Table D2), presumably due to outgassing of the titanium sample packet.

MDD models for samples WK-15, 7T-5, SR-17, and BN-11 approximately reproduced those samples’ measured diffusion and age data (Figs. D1, D2); however,

MDD models for samples PF-11 and SE-12 did not fit the experimental data well and were rejected. As a further test of the validity of the remaining models, we compared the age spectrum for each sample and its diffusion properties, quantified as log(r/r0), the ratio 234 between the apparent diffusion domain size and a reference domain size (Figure D2;

Richter et al., 1991). Plotted against cumulative 39Ar release, age and log(r/r0) should be highly correlated if laboratory argon loss proceeds by the same mechanism as natural argon loss over geological time periods, a key assumption of the MDD model (Lovera et al., 2002). Only steps which didn’t require Cl correction were included in calculation of the correlation coefficient, Cfg (except for sample 7T-5, for which all steps were Cl- corrected). Samples SR-17 and BN-11 showed very well correlated diffusion properties

(Cfg = 0.96–0.98), and sample WK-15 was moderately well correlated (Cfg = 0.85; Fig.

D2). Sample 7T-5 demonstrated a poor correlation between apparent age and diffusion properties (Cfg = 0.40 Fig. D2), but this may be strongly influenced by the fact that all the heating steps included in the correlation calculation yielded the same apparent age within error (75.6 ± 0.1 Ma, MSWD = 0.75), so the age spectrum being correlated was mostly noise. 235

TABLE D1. BIOTITE 40Ar/39Ar ANALYTICAL DATA

40 39 † 38 39 † 37 39 † 36 39 † 39 § 40 39 Sample ID Ar/ Ar Ar/ Ar Ar/ Ar Ar/ Ar ArK %40Ar* K/Ca Ar*/ ArK age, ± 1σ (10-15mol) (Ma)

Black Rock Range NBR1-09A-1 17.20 0.1702 0.0107 0.00443 1.51 92.2 196 15.86 105.6 ± 0.4 NBR1-09A-2 16.93 0.1596 0.0027 0.00423 2.29 92.4 786 15.65 104.2 ± 0.3 NBR1-09A-3 16.57 0.1670 0.0004 0.00209 2.81 96.1 4880 15.92 106.0 ± 0.2 NBR1-09A-4 16.14 0.1606 0.0005 0.00119 7.79 97.6 4083 15.76 104.9 ± 0.1 NBR1-09A-5 16.41 0.1597 0.0021 0.00175 1.90 96.7 1006 15.86 105.6 ± 0.3 NBR1-09A-6 16.72 0.1620 0.0016 0.00313 8.97 94.3 1291 15.77 105.0 ± 0.1 NBR1-09A-7 17.25 0.1628 0.0009 0.00437 8.16 92.4 2339 15.93 106.0 ± 0.1 J = 0.003799#; wt. mean = 105.30 Ma; 2σ std. error = 0.14 Ma; MSWD = 10.58 (2σ std. dev. = 1.3 Ma)

Seven Troughs Range 7T5-07A-1 16.01 0.0790 0.0583 0.00941 10.15 82.5 35 13.21 88.4 ± 0.2 7T5-07A-2 16.40 0.0964 0.0502 0.00999 7.59 81.9 41 13.43 89.9 ± 0.2 7T5-07A-3 17.24 0.0974 0.0501 0.01305 16.19 77.5 41 13.36 89.4 ± 0.2 7T5-07A-4 16.03 0.0888 0.0934 0.00898 16.37 83.3 21 13.35 89.3 ± 0.1 7T5-07A-5 15.59 0.0924 0.0488 0.00759 5.27 85.4 42 13.32 89.1 ± 0.2 7T5-07A-6 16.23 0.1003 0.0890 0.00968 13.24 82.2 23 13.35 89.3 ± 0.1 7T5-07A-7 15.90 0.1079 0.1040 0.00772 13.05 85.5 19 13.60 91.0 ± 0.1 7T5-07A-8 15.03 0.1020 0.0751 0.00606 25.00 87.9 27 13.21 88.5 ± 0.1 7T5-07A-9 14.87 0.0920 0.0581 0.00508 11.22 89.7 35 13.35 89.3 ± 0.1 J = 0.003803#; wt. mean = 89.30 Ma; 2σ std. error = 0.09 Ma; MSWD = 35.93 (2σ std. dev. = 1.5 Ma)

Bloody Run Hills BN11-08A-1 16.01 0.0560 0.0216 0.00673 8.32 87.4 96 13.99 93.5 ± 0.1 BN11-08A-2 15.48 0.0626 0.0092 0.00405 12.40 92.1 228 14.25 95.2 ± 0.1 BN11-08A-3 15.29 0.0548 0.0232 0.00440 7.45 91.3 89 13.96 93.4 ± 0.2 BN11-08A-4 14.66 0.0583 0.0228 0.00246 7.65 94.9 91 13.90 93.0 ± 0.1 BN11-08A-5 14.81 0.0629 0.0198 0.00222 7.90 95.4 105 14.13 94.4 ± 0.1 BN11-08A-6 15.11 0.0615 0.0091 0.00287 8.78 94.2 229 14.23 95.1 ± 0.1 BN11-08A-7 15.46 0.0601 0.0061 0.00401 6.97 92.1 346 14.24 95.2 ± 0.2 J = 0.003804#; wt. mean = 94.37 Ma; 2σ std. error = 0.10 Ma; MSWD = 51.61 (2σ std. dev. = 1.9 Ma)

*Radiogenic 40Ar (atmosphere corrected) †Ratios are background, deadtime (ion-collector measurements), decay, and detector-bias corrected §Sensitivity: 5.5×10-21 (mol/mV) #J-factor calculated from co-irradiated Taylor Creek sanidine, (28.34 ± 0.09 Ma; Renne et al., 1998) TABLE D2. K-FELDSPAR 40Ar/39Ar ANALYTICAL DATA

40 39 † 38 39 † 37 39 † 36 39 † 39 § 40 39 Temp. (°C) Duration Ar/ Ar Ar/ Ar Ar/ Ar Ar/ Ar ArK cumul. %40Ar* Cl/K Ar*/ ArK Age, ± 1σ Cl-corr. age, ± 1σ (min.) -15 39 (Ma) (10 mol) % ArK (Ma)

Wassuk Range: WK-15 300# 10 687.77 1.3749 1.2433 1.5029 0.02 0.01 35.44 2.9932 243.916 1230 ± 740 undefined 300# 17 4.87 0.7379 0.0312 0.0024 2.11 0.58 84.67 2.0074 4.123 29.7 ± 0.4 undefined 350# 14 218.40 0.5402 0.8008 0.3598 0.05 0.59 51.34 1.2744 112.184 672 ± 5 -3000 ± 15000 350# 18 157.06 0.4666 0.9098 0.4901 0.04 0.60 7.82 1.0025 12.292 87.0 ± 7.5 -5000 ± 34000 400 13 307.57 0.3038 0.1773 0.2696 0.30 0.69 74.10 0.6658 227.922 1174 ± 4 400 ± 1200 400 19 57.85 0.1590 0.3144 0.1612 0.13 0.72 17.64 0.3211 10.206 72.6 ± 2.4 -600 ± 1100 450 14 66.01 0.0902 0.0773 0.0777 0.69 0.91 65.17 0.1740 43.018 288 ± 1 10 ± 400 450 18 20.50 0.0402 0.0911 0.0580 0.44 1.03 16.27 0.0456 3.335 24.0 ± 0.8 -60 ± 110 500 15 24.61 0.0303 0.0403 0.0359 1.76 1.51 56.76 0.0298 13.972 98.6 ± 0.6 48 ± 67 500 18 11.04 0.0215 0.0587 0.0301 1.10 1.81 19.25 0.0085 2.124 15.3 ± 0.4 0 ± 20 533# 1 83.52 0.2246 24.0691 0.4477 0.00 1.81 -56.27 0.3605 -47.763 -390 ± 420 -1400 ± 2100 533 21 9.68 0.0177 0.0340 0.0257 1.97 2.34 21.22 0.0002 2.055 14.8 ± 0.3 14.4 ± 1.6 566 16 9.76 0.0178 0.0214 0.0243 4.95 3.69 26.05 0.0012 2.543 18.4 ± 0.3 16.2 ± 3.2 566 18 9.08 0.0168 0.0223 0.0232 2.79 4.45 24.24 0.0000 2.202 15.9 ± 0.5 15.9 ± 1.5 591 14 8.41 0.0162 0.0193 0.0201 3.55 5.42 28.90 0.0000 2.431 17.6 ± 0.3 17.6 ± 1.5 591 18 251.19 0.1694 0.0565 0.8457 0.69 5.61 0.50 0.0000 1.268 9.2 ± 8.3 9.2 ± 9.2 600 9 25.12 0.0273 0.0199 0.0744 3.64 6.60 12.39 0.0014 3.112 22.4 ± 0.8 19.9 ± 3.7 600 18 30.43 0.0299 0.0151 0.0956 5.14 8.00 7.03 0.0000 2.138 15.4 ± 1.0 not corrected 625 15 18.87 0.0226 0.0163 0.0559 5.15 9.41 12.35 0.0000 2.331 16.8 ± 0.6 not corrected 625 18 16.44 0.0211 0.0175 0.0476 4.08 10.52 14.28 0.0000 2.348 17.0 ± 0.5 not corrected 650 14 10.97 0.0164 0.0153 0.0293 4.76 11.82 20.76 0.0000 2.278 16.5 ± 0.4 not corrected 650 18 13.36 0.0190 0.0175 0.0365 4.26 12.98 19.07 0.0000 2.549 18.4 ± 0.4 not corrected 675 14 10.94 0.0173 0.0158 0.0276 5.56 14.49 25.24 0.0000 2.762 19.9 ± 0.3 not corrected 675** 18 16.26 0.0206 0.0175 0.0461 3.85 15.54 16.02 0.0000 2.606 18.8 ± 0.6 not corrected 700** 21 16.30 0.0206 0.0164 0.0461 4.71 16.83 16.32 0.0000 2.661 19.2 ± 0.5 not corrected 720** 14 18.96 0.0224 0.0192 0.0551 3.77 17.86 14.00 0.0000 2.655 19.2 ± 0.7 not corrected 740** 14 29.51 0.0291 0.0182 0.0914 4.77 19.16 8.40 0.0000 2.478 17.9 ± 1.1 not corrected fusion** 1 7.04 0.0133 0.0002 0.0020 296.55 100 91.38 0.0004 6.435 44.60 ± 0.09 not corrected J = 0.004021††; total gas age = 43.1 ± 0.3 Ma (2σ analytical precision)

*Radiogenic 40Ar (atmosphere corrected) †Ratios are baseline, deadtime (ion-collector measurements), decay, and detector-bias corrected §Sensitivity: 5.5×10-21 (mol/mV) #This step was not included in MDD modeling. **These steps were combined for MDD modeling ††J-factor calculated from co-irradiated Taylor Creek sanidine, (28.34 ± 0.09 Ma; Renne et al., 1998) 236 TABLE D2. K-FELDSPAR 40Ar/39Ar ANALYTICAL DATA (continued) Cl-corr. age, ± 1σ 40 39 † 38 39 † 37 39 † 36 39 † 39 § 39 40 39 Temp. (°C) Duration Ar/ Ar Ar/ Ar Ar/ Ar Ar/ Ar ArK % ArK %40Ar* Cl/K Ar*/ ArK age, ± 1σ (Ma) (Ma) (min.) Seven Troughs Range: 7T-5 300# 10 179.41 1.7497 1.9379 0.4179 0.02 0.01 31.23 4.5998 56.10 400 ± 2100 -300 ± 3500 300# 18 -50.90 -0.2044 -1.5149 -0.1711 -0.02 0.00 0.96 0.0000 -0.49 0 ± 810 0 ± 810 350# 14 191.67 1.8993 0.5984 0.3009 0.10 0.05 53.61 5.0711 102.80 620 ± 270 -50 ± 460 350# 18 256.89 2.1207 0.6269 0.4141 0.07 0.09 52.37 5.6257 134.59 780 ± 510 90 ± 830 400 14 139.89 1.0559 0.2331 0.1963 0.27 0.23 58.52 2.7877 81.88 510 ± 120 150 ± 170 400 22 28.59 0.1129 0.3438 0.0514 0.12 0.29 46.87 0.2505 13.40 95 ± 66 57 ± 69 450 17 41.87 0.2889 0.1094 0.0485 0.53 0.56 65.74 0.7395 27.53 189 ± 13 82 ± 20 450 22 18.06 0.0589 0.1477 0.0171 0.32 0.73 71.98 0.1187 13.00 92 ± 13 74 ± 14 450 29 17.64 0.0509 0.1523 0.0174 0.30 0.88 70.79 0.0965 12.49 88 ± 15 74 ± 16 450 44 21.71 0.0595 0.1248 0.0318 0.34 1.05 56.66 0.1129 12.30 87 ± 31 70 ± 32 500 17 26.70 0.1619 0.0648 0.0265 1.28 1.71 70.57 0.3992 18.84 131 ± 1 72.7 ± 7.0 500 22 13.20 0.0273 0.1255 0.0070 0.61 2.02 84.27 0.0365 11.12 78.8 ± 0.5 73.3 ± 0.8 500 29 12.78 0.0233 0.1373 0.0056 0.83 2.45 87.01 0.0263 11.12 78.7 ± 0.5 74.8 ± 0.7 500 41 15.31 0.0229 0.1530 0.0145 0.53 2.72 71.85 0.0203 11.00 78.0 ± 0.7 74.9 ± 0.8 550 17 12.60 0.0217 0.0965 0.0047 1.44 3.46 88.77 0.0223 11.18 79.2 ± 0.4 75.9 ± 0.6 550 22 11.74 0.0161 0.0711 0.0032 1.23 4.10 91.69 0.0074 10.77 76.3 ± 0.4 75.2 ± 0.4 550 29 11.83 0.0148 0.0600 0.0036 1.33 4.78 90.78 0.0036 10.74 76.1 ± 0.3 75.6 ± 0.4 550 41 12.04 0.0145 0.0606 0.0042 1.22 5.41 89.53 0.0026 10.78 76.4 ± 0.4 76.0 ± 0.4 575 17 11.74 0.0159 0.0642 0.0032 1.41 6.13 91.86 0.0068 10.79 76.5 ± 0.4 75.4 ± 0.4 575 22 11.55 0.0140 0.0656 0.0027 1.24 6.77 92.84 0.0019 10.72 76.0 ± 0.3 75.7 ± 0.3 600 17 11.74 0.0147 0.0607 0.0033 1.87 7.73 91.45 0.0036 10.74 76.1 ± 0.3 75.6 ± 0.4 600 22 11.40 0.0138 0.0560 0.0022 2.08 8.80 94.12 0.0016 10.73 76.0 ± 0.3 75.8 ± 0.3 625 17 12.76 0.0149 0.0584 0.0070 1.76 9.71 83.50 0.0021 10.66 75.5 ± 0.4 75.2 ± 0.4 625 22 12.44 0.0147 0.0581 0.0056 1.50 10.48 86.48 0.0024 10.76 76.2 ± 0.4 75.9 ± 0.4 650 17 11.67 0.0147 0.0577 0.0031 1.73 11.37 92.04 0.0036 10.74 76.1 ± 0.3 75.6 ± 0.4 650 22 11.72 0.0140 0.0564 0.0033 1.69 12.24 91.54 0.0017 10.73 76.1 ± 0.3 75.8 ± 0.4 675 17 11.84 0.0145 0.0563 0.0038 1.98 13.26 90.40 0.0027 10.70 75.9 ± 0.3 75.4 ± 0.4 675 22 12.19 0.0142 0.0559 0.0050 1.94 14.26 87.76 0.0012 10.70 75.8 ± 0.3 75.7 ± 0.4 700 17 16.11 0.0172 0.0571 0.0182 1.66 15.11 66.41 0.0028 10.70 75.9 ± 0.6 75.5 ± 0.6 700 22 18.77 0.0185 0.0616 0.0275 1.31 15.79 56.61 0.0016 10.62 75.3 ± 0.8 75.1 ± 0.8 fusion 1 13.18 0.0261 0.0005 0.0032 163.64 100.00 92.74 0.0351 12.22 86.4 ± 0.1 not corrected J = 0.004013††; total gas age = 86.2 ± 1.0 Ma (2σ analytical precision)

††J-factor calculated from co-irradiated Taylor Creek sanidine, (28.34 ± 0.09 Ma; Renne et al., 1998) 237 TABLE D2. K-FELDSPAR 40Ar/39Ar ANALYTICAL DATA (continued)

40 39 † 38 39 † 37 39 † 36 39 † 39 § 40 39 Temp. (°C) Duration Ar/ Ar Ar/ Ar Ar/ Ar Ar/ Ar ArK cumul. %40Ar* Cl/K Ar*/ ArK age, ± 1σ Cl-corr. age, ± 1σ (min.) -15 39 (Ma) (10 mol) % ArK (Ma)

Pine Forest Range: PF-11# 300 5 -109.06 -0.1236 -29.1194 -0.3408 0.00 0.00 9.69 0.0000 -10.36 -70 ± 910 -70 ± 990 400 10 212.08 0.2852 6.2873 0.3932 0.21 0.15 45.42 0.5524 96.73 565 ± 31 30 ± 88 400 22 66.64 0.0672 8.2049 0.1974 0.14 0.26 13.33 0.0486 8.93 60 ± 23 6 ± 27 450 17 38.99 0.0523 2.4105 0.0780 0.51 0.63 41.31 0.0691 16.13 107 ± 3 31 ± 10 450 22 34.48 0.0358 4.1983 0.0960 0.29 0.84 18.56 0.0140 6.42 43.5 ± 5.5 27.7 ± 6.4 500 23 22.20 0.0285 0.9908 0.0389 1.05 1.61 48.35 0.0233 10.74 72.2 ± 0.5 46.3 ± 3.3 500 23 24.85 0.0249 2.5754 0.0596 0.39 1.90 29.83 0.0028 7.43 50.2 ± 0.8 47.1 ± 1.4 500 23 31.92 0.0298 2.7376 0.0812 0.36 2.17 25.40 0.0051 8.12 54.8 ± 0.9 49.1 ± 1.6 500 41 42.73 0.0354 2.8404 0.1063 0.34 2.42 26.93 0.0075 11.53 77.4 ± 1.5 69.1 ± 2.1 550 18 19.90 0.0211 1.1112 0.0338 0.86 3.06 50.09 0.0053 9.97 67.1 ± 0.9 61.2 ± 1.5 550 18 24.33 0.0226 1.7510 0.0506 0.61 3.51 38.94 0.0010 9.49 63.9 ± 0.6 62.8 ± 1.2 550 29 29.69 0.0254 1.4119 0.0630 0.71 4.03 37.58 0.0021 11.17 75.0 ± 1.0 72.6 ± 1.5 550 41 35.83 0.0287 1.4084 0.0824 0.76 4.60 32.22 0.0011 11.56 77.5 ± 1.1 76.3 ± 1.5 575 18 27.81 0.0242 1.4193 0.0552 0.78 5.17 41.58 0.0028 11.57 77.7 ± 1.1 74.6 ± 1.5 575 18 32.48 0.0266 1.5650 0.0702 0.70 5.69 36.43 0.0017 11.84 79.4 ± 0.9 77.6 ± 1.4 575 29 38.17 0.0297 1.1827 0.0867 0.91 6.36 33.02 0.0017 12.61 84.5 ± 1.0 82.6 ± 1.5 575 41 42.80 0.0326 1.2330 0.1041 0.91 7.04 28.28 0.0008 12.11 81.2 ± 1.0 80.3 ± 1.4 600 18 34.35 0.0286 0.7024 0.0733 1.77 8.35 37.01 0.0057 12.72 85.2 ± 0.7 78.9 ± 1.4 600 23 69.31 0.0501 1.0530 0.1968 1.27 9.29 16.17 0.0010 11.22 75.3 ± 1.7 74.2 ± 2.1 600 29 101.04 0.0708 0.8111 0.3094 1.50 10.40 9.53 0.0000 9.64 64.9 ± 2.6 not corrected 600 41 122.78 0.0842 0.7123 0.3685 1.55 11.54 11.33 0.0059 13.92 93.0 ± 3.0 not corrected 625 11 2853.99 1.8269 8.7097 9.7734 0.13 11.64 -1.17 0.0000 -33.61 -250 ± 100 not corrected 625 22 282.07 0.1830 4.4941 0.9036 0.24 11.82 5.45 0.0014 15.41 103 ± 8 not corrected 625 29 190.26 0.1262 4.0380 0.5983 0.27 12.02 7.22 0.0028 13.76 92.0 ± 5.4 not corrected 625 41 165.59 0.1093 3.4821 0.5120 0.32 12.26 8.77 0.0010 14.56 97.1 ± 4.5 not corrected 650 18 202.15 0.1341 3.4791 0.6372 0.31 12.48 6.97 0.0044 14.12 94.3 ± 5.6 not corrected 650 23 178.49 0.1181 2.7541 0.5569 0.37 12.76 7.90 0.0020 14.12 94.3 ± 4.8 not corrected 650 29 172.72 0.1147 2.8828 0.5374 0.41 13.06 8.17 0.0028 14.15 94.5 ± 4.5 not corrected 650 41 176.60 0.1168 2.2217 0.5474 0.48 13.41 8.49 0.0032 15.01 100 ± 5 not corrected 675 18 257.04 0.1682 2.3921 0.8211 0.48 13.77 5.66 0.0033 14.58 97.3 ± 6.8 not corrected 675 23 229.49 0.1506 2.0297 0.7268 0.56 14.18 6.47 0.0036 14.86 99.1 ± 6.0 not corrected 700 18 301.34 0.1988 1.4110 0.9684 0.85 14.81 5.06 0.0113 15.25 102 ± 8 not corrected

Pine Forest Range: PF-11# (continued) 700 23 256.97 0.1694 1.2462 0.8175 0.96 15.52 6.02 0.0084 15.47 103 ± 7 not corrected 725 18 311.40 0.2075 1.0531 1.0290 1.22 16.42 2.37 0.0039 7.39 50.0 ± 8.4 not corrected 725 23 259.36 0.1728 0.9553 0.8423 1.30 17.38 4.05 0.0050 10.51 70.7 ± 6.9 not corrected 750 18 292.30 0.1949 0.7565 0.9708 1.71 18.65 1.87 0.0000 5.46 37.1 ± 8.0 not corrected 750 23 243.44 0.1631 0.6837 0.7998 1.80 19.98 2.92 0.0002 7.12 48.2 ± 6.6 not corrected fusion 1 13.72 0.0155 0.0005 0.0028 108.16 100.00 93.77 0.0059 12.86 86.1 ± 0.1 not corrected J = 0.003800††; total gas age = 84.5 ± 2.0 Ma (2σ analytical precision)

Sahwave Range: SE-12# 400 11 26.64 0.0329 6.2346 0.0517 0.23 0.21 44.33 0.0293 11.86 79.4 ± 0.8 65.6 ± 4.7 400 23 21.82 0.0254 3.0309 0.0378 0.46 0.61 49.72 0.0154 10.87 72.9 ± 0.5 65.6 ± 2.5 450 11 26.42 0.0264 11.7855 0.0573 0.11 0.71 39.19 0.0085 10.44 70.0 ± 1.1 66.0 ± 1.8 450 23 16.47 0.0168 1.8399 0.0198 0.73 1.36 65.21 0.0010 10.75 72.1 ± 0.3 71.6 ± 0.5 450 18 17.76 0.0172 2.9349 0.0241 0.42 1.73 60.98 0.0000 10.85 72.7 ± 0.4 not corrected 500 16 15.38 0.0157 1.7214 0.0158 0.74 2.39 70.25 0.0000 10.81 72.5 ± 0.3 not corrected 500 23 18.64 0.0176 2.8084 0.0266 0.45 2.80 58.79 0.0000 10.98 73.6 ± 0.4 not corrected 500 29 21.67 0.0195 3.1851 0.0368 0.40 3.15 50.75 0.0000 11.02 73.9 ± 0.5 not corrected 500 41 18.33 0.0172 1.6018 0.0251 0.79 3.85 60.03 0.0000 11.02 73.8 ± 0.4 not corrected 550 18 14.79 0.0150 1.6371 0.0135 0.76 4.52 73.69 0.0000 10.91 73.1 ± 0.3 not corrected 550 23 14.03 0.0145 1.0087 0.0104 1.22 5.60 78.44 0.0000 11.01 73.8 ± 0.2 not corrected 550 29 16.34 0.0160 1.5647 0.0184 0.82 6.33 67.30 0.0000 11.01 73.8 ± 0.3 not corrected 550 41 14.70 0.0148 0.7798 0.0125 1.64 7.78 75.03 0.0000 11.04 74.0 ± 0.2 not corrected 575 18 13.83 0.0143 1.1177 0.0098 1.13 8.78 79.53 0.0000 11.00 73.8 ± 0.2 not corrected 575 23 14.52 0.0148 1.1091 0.0120 1.14 9.79 75.88 0.0000 11.03 73.9 ± 0.2 not corrected 575 29 16.17 0.0157 1.3314 0.0173 0.94 10.62 68.76 0.0000 11.13 74.6 ± 0.3 not corrected 575 41 19.70 0.0178 1.6786 0.0288 0.73 11.27 57.34 0.0000 11.31 75.8 ± 0.4 not corrected 600 18 16.99 0.0163 2.3267 0.0202 0.51 11.72 65.65 0.0000 11.17 74.9 ± 0.4 not corrected 600 23 15.37 0.0152 1.2190 0.0145 1.07 12.66 72.51 0.0000 11.15 74.7 ± 0.3 not corrected 600 29 14.90 0.0148 0.8132 0.0128 1.78 14.24 74.93 0.0000 11.17 74.9 ± 0.3 not corrected 600 41 19.25 0.0176 1.1890 0.0270 1.04 15.16 58.93 0.0000 11.35 76.0 ± 0.4 not corrected 625 18 19.05 0.0183 2.3249 0.0272 0.63 15.72 58.55 0.0012 11.17 74.9 ± 0.4 not corrected

*Radiogenic 40Ar (atmosphere corrected) †Ratios are baseline, deadtime (ion-collector measurements), decay, and detector-bias corrected §Sensitivity: 5.5×10-21 (mol/mV) #This sample could not be adequately fit by MDD modeling †† J-factor calculated from co-irradiated Taylor Creek sanidine, (28.34 ± 0.09 Ma; Renne et al., 1998) 239 TABLE D2. K-FELDSPAR 40Ar/39Ar ANALYTICAL DATA (continued)

Sahwave Range: SE-12# (continued) 625 23 25.40 0.0220 2.4425 0.0462 0.55 16.21 46.89 0.0015 11.93 79.8 ± 0.6 not corrected 625 29 34.16 0.0273 3.6473 0.0741 0.35 16.52 36.63 0.0018 12.54 83.8 ± 0.9 not corrected 625 41 35.89 0.0286 3.1342 0.0798 0.42 16.90 34.88 0.0023 12.54 83.8 ± 0.9 not corrected 650 18 32.85 0.0273 4.3972 0.0697 0.31 17.17 38.21 0.0040 12.59 84.1 ± 0.9 not corrected 650 23 39.66 0.0312 4.0864 0.0901 0.30 17.44 33.61 0.0044 13.37 89.2 ± 1.1 not corrected 650 29 45.82 0.0351 5.2165 0.1102 0.24 17.65 29.73 0.0047 13.67 91.2 ± 1.3 not corrected 650 41 49.75 0.0371 4.6568 0.1230 0.26 17.88 27.62 0.0035 13.78 91.9 ± 1.4 not corrected 675 18 49.35 0.0377 5.6825 0.1165 0.23 18.08 31.08 0.0087 15.40 102 ± 1 not corrected 675 23 68.86 0.0481 4.9479 0.1705 0.26 18.31 27.32 0.0092 18.87 125 ± 2 not corrected 700 18 84.10 0.0572 4.0634 0.2121 0.34 18.61 25.81 0.0127 21.77 143 ± 2 not corrected 700 23 87.677698 0.0600801 3.6153399 0.2227276 0.38 18.95 25.21 0.0153 22.16 146 ± 2 not corrected 725 18 86.775015 0.0590902 3.0288574 0.2157005 0.52 19.41 26.78 0.0161 23.28 153 ± 2 not corrected 725 23 75.884205 0.0524825 2.6422376 0.1902274 0.50 19.85 26.15 0.0111 19.88 131 ± 2 not corrected 750 18 57.754565 0.0418034 2.78223 0.1361767 0.51 20.30 30.64 0.0096 17.73 117 ± 1 not corrected 750 23 49.349215 0.0372726 2.6612903 0.1157829 0.50 20.75 31.02 0.0077 15.33 102 ± 1 not corrected 775 18 40.394644 0.0320721 2.226483 0.0867055 0.70 21.37 36.92 0.0084 14.93 99.4 ± 1.0 not corrected 775 23 38.240764 0.0307227 2.0014685 0.0819058 0.77 22.05 37.03 0.0071 14.18 94.5 ± 0.9 not corrected 800 18 31.789304 0.0271377 1.5079623 0.0617383 1.23 23.14 42.88 0.0077 13.64 91.0 ± 0.7 not corrected 800 23 30.851129 0.0263475 1.1299769 0.0598543 1.47 24.44 42.85 0.0064 13.23 88.3 ± 0.6 not corrected fusion 1 12.409845 0.0145707 0.001228 0.0014508 85.28 100.00 96.31 0.0041 11.95 112 ± 8 not corrected J = 0.003792††; total gas age = 81.0 ± 0.4 Ma (2σ analytical precision)

Santa Rosa Range: SR-17 400 11 44.51 0.0621 8.5470 0.1161 0.17 0.22 24.30 0.0770 10.88 72.9 ± 1.4 20 ± 12 400 23 32.12 0.0407 8.0132 0.0915 0.15 0.43 17.62 0.0303 5.689 38.5 ± 1.4 17.4 ± 5.1 450 18 10.01 0.0206 2.2308 0.0235 0.58 1.19 32.00 0.0094 3.207 21.8 ± 0.3 15.2 ± 1.7 450 23 10.54 0.0189 2.8300 0.0271 0.43 1.76 25.81 0.0031 2.727 18.6 ± 0.4 16.4 ± 0.9 500 18 6.26 0.0152 1.0998 0.0103 1.23 3.40 52.38 0.0013 3.282 22.3 ± 0.2 21.4 ± 0.7 500 23 6.84 0.0153 1.6763 0.0140 0.79 4.45 41.05 0.0000 2.812 19.1 ± 0.2 19.1 ± 0.7 500 29 8.88 0.0163 2.1623 0.0199 0.62 5.28 35.40 0.0000 3.149 21.4 ± 0.3 21.4 ± 0.7 500 41 11.65 0.0179 2.3509 0.0281 0.55 6.01 30.05 0.0000 3.507 23.8 ± 0.4 23.8 ± 0.7

*Radiogenic 40Ar (atmosphere corrected) †Ratios are baseline, deadtime (ion-collector measurements), decay, and detector-bias corrected §Sensitivity: 5.5×10-21 (mol/mV) #This sample could not be adequately fit by MDD modeling †† J-factor calculated from co-irradiated Taylor Creek sanidine, (28.34 ± 0.09 Ma; Renne et al., 1998) 240 TABLE D2. K-FELDSPAR 40Ar/39Ar ANALYTICAL DATA (continued)

Santa Rosa Range: SR-17 (continued) 550 18 7.29 0.0149 1.2743 0.0098 1.04 7.39 61.02 0.0008 4.452 30.2 ± 0.2 29.6 ± 0.7 550 23 8.20 0.0150 1.7556 0.0135 0.75 8.40 52.46 0.0000 4.308 29.2 ± 0.2 29.2 ± 0.7 550 29 10.53 0.0164 2.1860 0.0198 0.60 9.21 45.82 0.0000 4.832 32.7 ± 0.3 32.7 ± 0.7 550 41 13.58 0.0181 2.5032 0.0291 0.53 9.91 37.87 0.0000 5.152 34.9 ± 0.4 34.9 ± 0.7 575 18 10.70 0.0163 3.0552 0.0185 0.42 10.47 50.69 0.0001 5.436 36.8 ± 0.4 36.8 ± 0.7 575 23 12.33 0.0171 3.2860 0.0237 0.40 11.00 45.03 0.0000 5.566 37.7 ± 0.4 not corrected 575 29 15.49 0.0190 3.3980 0.0334 0.37 11.48 37.64 0.0000 5.842 39.5 ± 0.5 not corrected 575 41 17.96 0.0204 3.5419 0.0417 0.36 11.96 32.78 0.0000 5.900 39.9 ± 0.6 not corrected 600 18 13.33 0.0172 4.1623 0.0256 0.30 12.36 45.34 0.0000 6.061 41.0 ± 0.5 not corrected 600 23 14.89 0.0186 4.0270 0.0308 0.30 12.76 40.62 0.0001 6.064 41.0 ± 0.5 not corrected 600 29 17.22 0.0200 4.2927 0.0383 0.29 13.15 35.91 0.0002 6.201 41.9 ± 0.6 not corrected 600 41 24.60 0.0245 4.1067 0.0624 0.30 13.55 26.16 0.0002 6.452 43.6 ± 0.8 not corrected 625# 18 15.08 0.0187 4.9905 0.0309 0.25 13.88 41.69 0.0006 6.307 42.6 ± 0.6 not corrected 625# 23 16.47 0.0196 4.6302 0.0353 0.26 14.23 38.62 0.0006 6.380 43.1 ± 0.6 not corrected 625# 29 18.77 0.0209 4.6893 0.0438 0.26 14.57 32.77 0.0000 6.168 41.7 ± 0.7 not corrected 625# 41 22.25 0.0231 4.6540 0.0545 0.27 14.93 29.07 0.0005 6.488 43.8 ± 0.7 not corrected 650# 18 16.13 0.0195 5.3143 0.0353 0.22 15.23 37.64 0.0003 6.094 41.2 ± 0.6 not corrected 650# 23 17.47 0.0204 5.1582 0.0399 0.24 15.56 34.59 0.0006 6.063 41.0 ± 0.6 not corrected 650# 29 19.66 0.0216 5.0245 0.0469 0.25 15.88 31.23 0.0003 6.163 41.7 ± 0.7 not corrected 650# 41 22.76 0.0233 4.6942 0.0569 0.26 16.23 27.62 0.0000 6.305 42.6 ± 0.8 not corrected 675# 18 16.72 0.0200 5.9108 0.0377 0.22 16.52 35.79 0.0007 6.008 40.6 ± 0.6 not corrected 675# 23 18.39 0.0215 5.3221 0.0435 0.23 16.83 32.19 0.0017 5.943 40.2 ± 0.7 not corrected 700# 18 14.60 0.0184 4.1832 0.0311 0.30 17.23 38.97 0.0000 5.704 38.6 ± 0.5 not corrected 700# 23 15.40 0.0192 3.9322 0.0337 0.31 17.64 37.13 0.0004 5.735 38.8 ± 0.5 not corrected 725# 18 13.10 0.0177 3.3726 0.0262 0.38 18.15 42.72 0.0000 5.608 38.0 ± 0.4 not corrected 725# 23 13.36 0.0179 3.0420 0.0270 0.43 18.72 41.72 0.0001 5.585 37.8 ± 0.4 not corrected 750# 18 10.69 0.0160 1.4925 0.0178 0.91 19.94 51.67 0.0000 5.527 37.4 ± 0.3 not corrected 750# 23 11.03 0.0164 1.7770 0.0193 0.72 20.90 49.24 0.0000 5.438 36.8 ± 0.3 not corrected 775# 18 10.46 0.0161 1.5242 0.0170 0.85 22.04 52.67 0.0003 5.514 37.3 ± 0.3 not corrected 775# 23 11.17 0.0164 1.8531 0.0195 0.69 22.96 49.48 0.0000 5.533 37.5 ± 0.3 not corrected 800# 18 10.65 0.0162 1.6201 0.0172 0.81 24.04 53.23 0.0007 5.674 38.4 ± 0.3 not corrected 800# 23 11.41 0.0165 1.7525 0.0195 0.73 25.01 50.48 0.0002 5.766 39.0 ± 0.3 not corrected

Santa Rosa Range: SR-17 (continued) 825# 18 11.24 0.0167 1.5327 0.0182 0.91 26.22 52.87 0.0015 5.950 40.2 ± 0.3 not corrected 825# 23 12.03 0.0170 1.4767 0.0206 0.93 27.46 50.08 0.0009 6.028 40.8 ± 0.3 not corrected 850# 18 12.00 0.0172 1.4153 0.0204 1.08 28.90 50.34 0.0017 6.046 40.9 ± 0.3 not corrected 850# 23 12.20 0.0170 1.3249 0.0200 1.04 30.29 52.24 0.0014 6.382 43.1 ± 0.3 not corrected 875# 18 12.16 0.0172 1.2941 0.0198 1.16 31.84 52.44 0.0020 6.380 43.1 ± 0.3 not corrected 875# 23 12.84 0.0177 1.1880 0.0214 1.18 33.41 51.30 0.0026 6.590 44.5 ± 0.3 not corrected 900# 18 13.31 0.0182 1.2367 0.0220 1.24 35.06 51.58 0.0035 6.869 46.4 ± 0.3 not corrected fusion# 1 8.83 0.0147 0.0005 0.0025 48.66 100.00 91.38 0.0041 8.068 54.35 ± 0.08 not corrected J = 0.003791††; total gas age = 48.2 ± 0.4 Ma (2σ analytical precision)

Bloody Run Hills: BN-11 400 11 65.21 0.1158 0.6480 0.1918 0.25 0.10 13.09 0.1857 8.542 84.1 ± 3.0 34.9 ± 7.9 400 23 62.73 0.0843 0.4298 0.1839 0.18 0.17 13.40 0.1024 8.405 82.8 ± 2.9 55.8 ± 4.9 450 18 25.12 0.0510 0.2614 0.0561 0.61 0.41 33.99 0.0765 8.541 84.1 ± 0.9 64.0 ± 3.1 450 23 19.97 0.0299 0.2072 0.0435 0.44 0.59 35.63 0.0246 7.117 70.4 ± 0.8 63.9 ± 1.3 500 17 14.01 0.0259 0.2139 0.0213 1.57 1.21 54.89 0.0252 7.690 75.9 ± 0.4 69.3 ± 1.1 500 23 10.91 0.0173 0.2352 0.0130 1.15 1.67 64.55 0.0056 7.043 69.7 ± 0.3 68.2 ± 0.4 500 29 11.90 0.0173 0.2337 0.0161 0.99 2.07 59.95 0.0042 7.138 70.6 ± 0.4 69.5 ± 0.5 500 41 12.82 0.0176 0.2237 0.0188 0.97 2.45 56.65 0.0036 7.267 71.8 ± 0.4 70.9 ± 0.5 550 18 11.57 0.0192 0.2696 0.0135 2.40 3.41 65.37 0.0108 7.566 74.7 ± 0.3 71.9 ± 0.6 550 23 9.65 0.0148 0.2688 0.0076 2.06 4.23 76.77 0.0016 7.411 73.2 ± 0.2 72.8 ± 0.3 550 29 9.93 0.0149 0.2724 0.0082 1.91 4.99 75.63 0.0016 7.508 74.2 ± 0.2 73.7 ± 0.3 550 41 10.42 0.0148 0.2765 0.0096 1.87 5.74 72.70 0.0006 7.579 74.8 ± 0.3 74.7 ± 0.3 575 18 10.35 0.0158 0.3353 0.0091 1.73 6.43 74.07 0.0037 7.669 75.7 ± 0.3 74.7 ± 0.4 575 23 9.68 0.0146 0.3502 0.0069 1.77 7.13 78.78 0.0014 7.625 75.3 ± 0.2 74.9 ± 0.3 575 29 9.93 0.0145 0.3468 0.0076 1.83 7.86 77.25 0.0007 7.671 75.7 ± 0.2 75.6 ± 0.3 575 41 10.22 0.0145 0.3384 0.0084 2.03 8.67 75.77 0.0003 7.746 76.5 ± 0.3 76.4 ± 0.3 600 18 10.22 0.0157 0.3861 0.0080 2.06 9.50 76.73 0.0038 7.843 77.4 ± 0.3 76.4 ± 0.4 600 23 9.35 0.0137 0.3607 0.0054 2.17 10.36 82.79 0.0000 7.743 76.4 ± 0.2 76.4 ± 0.3 600 29 9.56 0.0138 0.3332 0.0061 2.29 11.27 81.22 0.0000 7.765 76.6 ± 0.2 76.6 ± 0.3 600 41 9.98 0.0139 0.3108 0.0073 2.54 12.29 78.38 0.0000 7.824 77.2 ± 0.2 77.2 ± 0.3

*Radiogenic 40Ar (atmosphere corrected) †Ratios are baseline, deadtime (ion-collector measurements), decay, and detector-bias corrected §Sensitivity: 5.5×10-21 (mol/mV) #These steps were combined for MDD modeling †† J-factor calculated from co-irradiated Taylor Creek sanidine, (28.34 ± 0.09 Ma; Renne et al., 1998) 242 TABLE D2. K-FELDSPAR 40Ar/39Ar ANALYTICAL DATA (continued)

Bloody Run Hills: BN-11 625 18 9.94 0.0150 0.3344 0.0069 2.39 13.24 79.33 0.0025 7.884 77.8 ± 0.2 77.1 ± 0.3 625 23 9.30 0.0136 0.3144 0.0049 2.55 14.26 84.27 0.0000 7.841 77.4 ± 0.2 77.4 ± 0.3 625 29 9.38 0.0134 0.2893 0.0051 2.69 15.33 83.95 0.0000 7.879 77.7 ± 0.2 77.7 ± 0.3 625 41 9.60 0.0136 0.2677 0.0057 3.01 16.53 82.46 0.0000 7.920 78.1 ± 0.2 78.1 ± 0.3 650 18 10.36 0.0153 0.2857 0.0080 2.75 17.63 76.99 0.0028 7.975 78.7 ± 0.2 77.9 ± 0.4 650 23 9.11 0.0133 0.2670 0.0040 2.91 18.79 86.94 0.0000 7.921 78.2 ± 0.2 78.2 ± 0.3 650 29 9.24 0.0134 0.2476 0.0043 3.10 20.02 86.15 0.0000 7.966 78.6 ± 0.2 78.6 ± 0.3 650 41 9.83 0.0138 0.2302 0.0061 3.34 21.35 81.42 0.0000 8.002 78.9 ± 0.2 78.9 ± 0.3 675 18 9.21 0.0136 0.2484 0.0041 2.93 22.52 86.70 0.0000 7.988 78.8 ± 0.2 78.8 ± 0.3 675 23 9.15 0.0133 0.2368 0.0039 3.25 23.82 87.36 0.0000 7.993 78.9 ± 0.2 78.9 ± 0.3 700 18 9.38 0.0140 0.2542 0.0046 4.54 25.63 85.54 0.0010 8.021 79.1 ± 0.2 78.8 ± 0.3 700 23 8.82 0.0129 0.2290 0.0027 4.62 27.47 90.79 0.0000 8.011 79.0 ± 0.2 79.0 ± 0.3 725 18 9.12 0.0135 0.2316 0.0036 5.90 29.82 88.23 0.0000 8.044 79.3 ± 0.2 79.3 ± 0.3 725 23 8.82 0.0129 0.2059 0.0026 5.80 32.14 91.25 0.0000 8.047 79.4 ± 0.2 not corrected 750 18 9.21 0.0138 0.2168 0.0036 4.15 33.79 88.18 0.0008 8.127 80.1 ± 0.2 not corrected 750 22 8.99 0.0132 0.1831 0.0028 3.69 35.26 90.78 0.0000 8.164 80.5 ± 0.2 not corrected 775 17 9.64 0.0144 0.1923 0.0049 4.08 36.89 84.92 0.0018 8.183 80.7 ± 0.2 not corrected 775 22 9.07 0.0132 0.1541 0.0028 3.82 38.41 90.58 0.0000 8.216 81.0 ± 0.2 not corrected 800 17 9.92 0.0148 0.1823 0.0057 4.82 40.33 82.82 0.0025 8.215 81.0 ± 0.2 not corrected 800 22 9.21 0.0134 0.1332 0.0033 3.75 41.82 89.28 0.0000 8.226 81.1 ± 0.2 not corrected 825 17 9.88 0.0149 0.1771 0.0054 3.75 43.32 83.70 0.0029 8.267 81.5 ± 0.3 not corrected 825 22 9.89 0.0140 0.1264 0.0054 3.55 44.74 83.77 0.0004 8.283 81.6 ± 0.3 not corrected 850 17 9.61 0.0141 0.1331 0.0043 3.59 46.17 86.58 0.0015 8.317 82.0 ± 0.3 not corrected 850 22 9.61 0.0139 0.1167 0.0044 3.31 47.49 86.17 0.0007 8.280 81.6 ± 0.3 not corrected 875 17 9.97 0.0147 0.1294 0.0053 3.41 48.85 83.99 0.0025 8.375 82.5 ± 0.3 not corrected 875 22 10.05 0.0146 0.1282 0.0056 3.67 50.31 83.30 0.0021 8.377 82.5 ± 0.3 not corrected 900 17 10.52 0.0155 0.1537 0.0071 3.14 51.56 79.76 0.0038 8.395 82.7 ± 0.3 not corrected fusion 1 10.38 0.0180 0.2165 0.0030 121.51 100.00 91.46 0.0128 9.472 93.1 ± 0.1 not corrected J = 0.005589††; total gas age = 85.8 ± 0.4 Ma (2σ analytical precision)

-1 1/T (K ) 1/T (K-1 ) 8 9 10 11 12 13 14 15 16 8 9 10 11 12 13 14 15 16 -4 -4

-5 Step-heating data -5 Model results -6 -6 2

-7 -7

log(D/r ) -8 WK-15 -8 7T-5 E = 46.6 ± 2.6 kcal/mol E = 48.6 ± 3.9 kcal/mol -9 2 -9 2 log(D0/r0 ) = 5.9 ± 0.8 log(D0/r0 ) = 6.8 ± 1.2 -10 -10

8 9 10 11 12 13 14 15 16 8 9 10 11 12 13 14 15 16 -4 -4

-5 -5

-6

2 -6

-7 -7

log(D/r ) -8 SR-17 -8 BN-11

E = 47.2 ± 3.0 kcal/mol -9 E = 42.2 ± 1.5 kcal/mol -9 2 2 log(D0/r0 ) = 7.2 ± 0.9 log(D0/r0 ) = 4.8 ± 0.4 -10 -10

Figure D1. Arrhenius plots showing measured log(D0/r0 2 ) and model fits for sucessfully modeled samples. All plots at same scale. 245

25 80 WK-15 79 7T-5 20 78 77 15 76 75 Stepheating results 10 74 (1σ range) Cl-corrected results 73

Apparent age (Ma) 5 72 Model results 71 0 70

0.80 0.80

) 0.60 0.60 0

0.40 0.40 log(r/r 0.20 0.20

0.00 0.00 Cfg = 0.85 Cfg = 0.40 -0.20 -0.20 1 3 5 7 9 11 13 15 0 2 4 6 8 10 12 14 16 18 20 55 90

50 SR-17 85 BN-11 45 80 40 75 35 30 70

25 65 20 60

Apparent age (Ma) 15 55 10 5 50

0.80 0.80

) 0.60 0.60 0

0.40 0.40 log(r/r 0.20 0.20

0.00 0.00 Cfg = 0.98 Cfg = 0.96 -0.20 -0.20 1 3 5 7 9 11 13 15 0 10 20 30 40 50 60 39 39 cumulative ArK release (%) cumulative ArK release (%)

Figure D2. Apparent age (top plots) and log(r/r0) spectra (bottom plots); same scale for each pair. Scales differ between samples. Measured age spectra are compared against model results; models and data are difficult to discern for samples SR-17 and BN-11 due to close agreement. Correlated part of the log(r/r0) spectrum shown in bold (see text). 246

APPENDIX E: U-Pb SHRIMP ANALYTICAL PROCEDURES

Sample zircons and chips of R-33 standard zircon were picked into 1 by 10 mm rows on double-stick tape on glass and then mounted in epoxy, which was polished down until the grains were exposed about halfway through their depth. The mounted grains were washed in dilute hydrochloric acid and distilled water. Optical microphotographs were taken of the zircons under reflected light, and cathodoluminescence imaging was performed using a JEOL 5600 scanning electron microscope. Images were used to reveal zonation and avoid cracks, inclusions, and other physical defects. Samples were gold- coated and dried in a vacuum oven. The Stanford/U.S.G.S. SHRIMP-RG (sensitive, high resolution ion microprobe – reverse geometry) was used to measure U, Th, and Pb

- isotopes as well as Zr, Hf, La, Ce, Nd, Sm, Eu, Gd, Dy, Er, and Yb using an O2 primary ion beam between 4–6 nA and a spot size of 20–30 µm. The primary ion beam was rastered over each spot for ~120 seconds to remove contamination before isotope analysis in three to five scans of the magnet. Isotope ratios were normalized using CZ3 concentration standard and zircon age standard R-33 (419 Ma; Black et al., 2004).

SQUID and ISOPLOT software (Ludwig, 2001, 2003) were used to reduce the data, yielding 207Pb corrected 206Pb/238Pb weighted-average ages (Figure 4.6; Table E1).

247

TABLE E1. U-Pb SHRIMP ANALYTICAL DATA

Spot number Common 206Pb U Th 232Th/238U 206Pb/238U Age ± 1σ Total 238U/206Pb (± Total 207Pb/206Pb (± (%) (ppm) (ppm) (Ma) % err) % err)

NVB-287: Seven Troughs Range NVB287-7 2.81 191 118 0.64 97.9 ± 1.6 63.51 ± 1.5 0.0703 ± 6.0 NVB287-8 0.57 161 74 0.48 105.5 ± 1.6 60.23 ± 1.5 0.0527 ± 5.4 NVB287-9 0.56 144 59 0.42 107.3 ± 1.7 59.26 ± 1.6 0.0526 ± 5.6 NVB287-10 0.68 123 94 0.79 104.8 ± 1.8 60.62 ± 1.7 0.0535 ± 6.8 NVB287-11 0.46 153 71 0.48 103.9 ± 1.6 61.27 ± 1.6 0.0518 ± 5.5 NVB287-12 0.84 128 84 0.68 103.6 ± 1.8 61.23 ± 1.7 0.0548 ± 6.0 NVB287-13 -0.02 167 65 0.40 104.3 ±1.6 61.30 ± 1.5 0.0480 ± 5.4 NVB287-14 0.22 145 64 0.45 102.9 ± 1.6 62.00 ± 1.6 0.0498 ± 5.6 NVB287-15 0.22 172 83 0.50 106.4 ± 1.5 59.95 ± 1.4 0.0499 ± 5.1 NVB287-16 0.32 233 170 0.75 101.8 ± 1.3 62.61 ± 1.2 0.0506 ± 4.5 NVB287-17 0.24 215 94 0.45 99.7 ± 1.3 63.99 ± 1.3 0.0499 ± 4.8 NVB287-18 0.55 658 676 1.06 101.2 ± 0.8 62.84 ± 0.7 0.0524 ± 2.6 NVB287-19 0.26 223 162 0.75 100.5 ± 1.3 63.46 ± 1.3 0.0501 ± 4.6 NVB287-20 0.20 389 174 0.46 103.6 ± 1.0 61.58 ± 0.9 0.0497 ± 3.4

DD-8: Diamond Mountains DD8-1 -0.06 206 76 0.38 109.1 ± 1.4 58.63 ± 1.2 0.0477 ± 4.7 DD8-2 1.29 130 61 0.48 108.5 ± 1.7 58.14 ± 1.5 0.0585 ± 5.2 DD8-3 0.69 99 56 0.58 108.8 ± 2.0 58.34 ± 1.7 0.0536 ± 7.6 DD8-4 0.41 113 59 0.54 105.6 ± 1.7 60.29 ± 1.6 0.0514 ± 6.1 DD8-5 0.11 254 106 0.43 104.0 ± 1.2 61.40 ± 1.1 0.0490 ± 4.3 DD8-6 -0.09 362 125 0.36 107.6 ± 1.1 59.44 ± 1.0 0.0475 ± 3.7 DD8-7 0.25 243 150 0.64 106.1 ± 1.3 60.11 ± 1.2 0.0501 ± 4.3 DD8-8 -0.15 218 147 0.70 110.5 ± 1.3 57.92 ± 1.1 0.0471 ± 4.4 DD8-9 -0.11 187 113 0.62 106.0 ± 1.4 60.38 ± 1.3 0.0473 ± 5.8 DD8-10 0.83 279 94 0.35 103.6 ± 1.1 61.19 ± 1.1 0.0546 ± 3.9 DD8-11 -0.03 344 118 0.35 109.8 ± 1.0 58.22 ± 1.1 0.0480 ± 3.4 DD8-12 0.02 231 95 0.42 108.9 ± 1.2 58.68 ± 1.1 0.0484 ± 4.2 DD8-13 -0.13 103 45 0.45 102.3 ± 1.9 62.57 ± 1.9 0.0470 ± 7.2 DD8-14 -0.36 219 73 0.34 109.7 ± 1.3 58.47 ± 1.2 0.0453 ± 4.7 DD8-15 0.41 96 40 0.43 108.8 ± 2.0 58.53 ± 1.8 0.0515 ± 6.7 DD8-16 0.22 166 103 0.64 105.8 ± 1.5 60.32 ± 1.4 0.0499 ± 5.2 DD8-17 0.43 180 78 0.45 102.9 ± 1.4 61.90 ± 1.4 0.0515 ± 5.1 DD8-18 -0.21 141 61 0.45 104.9 ± 1.6 61.08 ± 1.5 0.0464 ± 5.8 DD8-19 0.80 243 125 0.53 103.4 ± 1.3 61.35 ± 1.2 0.0544 ± 4.3 DD8-20 0.64 120 49 0.42 104.5 ± 1.6 60.83 ± 1.5 0.0532 ± 5.4 DD8-21 -0.07 117 52 0.45 104.1 ± 1.7 61.49 ± 1.6 0.0475 ± 6.0 DD8-22 -0.33 84 46 0.56 107.2 ± 1.9 59.81 ± 1.7 0.0455 ± 6.6

SL-11: Selenite Range (northern) SL11-1 -0.21 564 209 0.38 97.4 ± 0.7 65.79 ± 0.7 0.0463 ± 2.9 SL11-2 0.46 361 127 0.36 95.7 ± 0.9 66.55 ± 0.9 0.0516 ± 3.6 SL11-3 0.01 428 137 0.33 98.6 ± 0.9 64.87 ± 0.9 0.0481 ± 3.5 SL11-4 0.14 460 162 0.36 96.9 ± 0.8 65.94 ± 0.8 0.0491 ± 3.2 SL11-5 0.07 672 70 0.11 95.0 ± 0.8 67.29 ± 0.9 0.0485 ± 2.8 SL11-6 -0.31 339 107 0.33 96.3 ± 1.0 66.62 ± 1.0 0.0455 ± 4.2 SL11-7 0.14 285 84 0.30 97.1 ± 1.1 65.82 ± 1.1 0.0491 ± 4.2 SL11-8 0.20 640 105 0.17 95.1 ± 0.7 67.16 ± 0.7 0.0495 ± 2.7 SL11-9 0.25 181 68 0.39 96.7 ± 1.4 66.03 ± 1.4 0.0499 ± 5.5 SL11-10 0.25 283 82 0.30 93.9 ± 1.1 67.96 ± 1.1 0.0499 ± 4.2