Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada (GSA Field Guide 11)

. THE GEOLOGICAL SOCIETY • OF AMERICA® Field Guide 11 H~ ~ b~i~~ t~ r~t~n~. V~ ~~n~~~. ~~~ t~. ~~~f~. mn~~~~r~. ~n~ ...

Author: Ernest M. Duebendorfer | Eugene Irwin Smith

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. . THE GEOLOGICAL SOCIETY • OF AMERICA®

Field Guide 11

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edited by Ernest M. Duebendorter and Eugene I. Smith

Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada

edited by Ernest M. Duebendorfer Northern Arizona University Geology Department Frier Hall Knoles Drive Flagstaff, Arizona 86011-4099 USA Eugene I. Smith Department of Geoscience University of Nevada, Las Vegas 4505 S. Maryland Parkway Las Vegas, Nevada 89154-4010 USA

Field Guide 11 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140 USA

2008

Copyright © 2008, The Geological Society of America, Inc. (GSA). All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact the Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. Library of Congress Cataloging-in-Publication Data Field guide to plutons, volcanoes, faults, reefs, dinosaurs, and possible glaciation in selectedareas of Arizona, California, and Nevada / edited by Ernest M. Duebendorfer, Eugene I. Smith. p. cm. -- (Field guide ; 11) Includes bibliographical references. ISBN: 978-0-8137-0011-3 (pbk.) 1. Geology--Arizona. 2. Geology--California. 3. Geology--Nevada. I. Duebendorfer, Ernest M. II. Smith, Eugene I. (Eugene Irwin), 1944QE85.F54 2008 557.9--dc22 2008006898 Cover: Spectacular geology in the Lake Mead area just west of Las Vegas. The River Mountains volcanic section (foreground in Nevada) and the Wilson Ridge pluton (on the skyline to the east in Arizona) represent a linked volcanic-plutonic system separated by the Saddle Island detachment fault. The mesa is Fortification Hill capped by 5.8 m.y. old basalt. Photo by Eugene Smith, May 2006.

10 9 8 7 6 5 4 3 2 1

ii

Contents

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section . . . . . . . . . . . . 1 Denise Honn and Eugene I. Smith 2. Late Paleozoic deformation in central and southern Nevada . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 Pat Cashman, Jim Trexler, Walt Snyder, Vladimir Davydov, and Wanda Taylor 3. Active tectonics of the eastern California shear zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 Kurt L. Frankel, Allen F. Glazner, Eric Kirby, Francis C. Monastero, Michael D. Strane, Michael E. Oskin, Jeffrey R. Unruh, J. Douglas Walker, Sridhar Anandakrishnan, John M. Bartley, Drew S. Coleman, James F. Dolan, Robert C. Finkel, Dave Greene, Andrew Kylander-Clark, Shasta Marrero, Lewis A. Owen, and Fred Phillips 4. Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic-Paleozoic boundary . . . . 83 Stephen M. Rowland, Lynn K. Oliver, and Melissa Hicks 5. Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 Andrew P. Barth, J. Lawford Anderson, Carl E. Jacobson, Scott R. Paterson, and Joseph L. Wooden 6. Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119 James E. Faulds, Keith A. Howard, and Ernest M. Duebendorfer 7. Interpretation of Pleistocene glaciation in the Spring Mountains of Nevada: Pros and Cons . . . 153 Jerry Osborn, Matthew Lachniet, and Marvin (Nick) Saines 8. Quaternary volcanism in the San Francisco Volcanic Field: Recent basaltic eruptions that profoundly impacted the northern Arizona landscape and disrupted the lives of nearby residents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 S.L. Hanson, W. Duffield, and J. Plescia 9. The Spirit Mountain batholith and Secret Pass Canyon volcanic center: A cross-sectional view of the magmatic architecture of the uppermost crust of an extensional terrain, Colorado River, Nevada-Arizona . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 187 Nicholas P. Lang, B.J. Walker, Lily L. Claiborne, Calvin F. Miller, Richard W. Hazlett, and Matthew T. Heizler iii

iv

Contents

10. Devonian carbonate platform of eastern Nevada: Facies, surfaces, cycles, sequences, reefs, and catastrophic Alamo Impact Breccia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 215 John E. Warme, Jared R. Morrow, and Charles A. Sandberg 11. Dinosaurs and dunes! Sedimentology and paleontology of the Mesozoic in the Valley of Fire State Park . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249 Joshua W. Bonde, David J. Varricchio, Frankie D. Jackson, David B. Loope, and Aubrey M. Shirk

Preface

Welcome to Las Vegas! This guidebook has been prepared in conjunction with the 2008 combined Cordilleran and Rocky Mountain Sections meeting of the Geological Society of America. This volume contains background information and road logs for eleven field trips in Nevada, Arizona, and California. Southern Nevada and adjoining areas contain a rich geologic history spanning the interval from the Paleoproterozoic to the present. Las Vegas lies at or near several critical geological junctures and localities including the structural boundary between the Colorado Plateau and Basin and Range, the physiographic boundary between the Great Basin and the southern Basin and Range, the eastern margin of the Sevier foldand-thrust belt, the tectonically active Death Valley area, tilted and faulted volcanic-plutonic systems exposing the upper part of the crust, and the enigmatic “amagmatic zone.” Field trips in this volume span the geologic record from the Ediacaran (late Neoproterozoic) to the Holocene. Steve Rowland, Lynn Oliver, and Melissa Hicks will lead participants to three of the best examples of Ediacaran and Early Cambrian reefs in North America (Chapter 4). A trip led by John Warme, Jared Morrow, and Charles Sandberg (Chapter 10) examines the long-lived Devonian shallow-water carbonate platform and features a visit to the spectacular Alamo Impact Breccia. Middle Mississippian to late Permian tectonism as recorded by regional unconformities, folding, thrusting, and the stratigraphic record is the focus of a trip led by Pat Cashman, Jim Trexler, Walt Snyder, Vladimir Davydov, and Wanda Taylor (Chapter 2). Andy Barth, Lawford Anderson, Carl Jacobson, Scott Paterson, and Joe Wooden bring us into the Mesozoic with an overview of the tectonic evolution of a tilted section through the upper and middle crust of the Cretaceous Cordilleran arc (Chapter 5). Cretaceous sedimentary rocks deposited in the foredeep of the Sevier fold-and-thrust belt and their dinosaur fossils are the topic of a trip led by Joshua Bonde, David Varricchio, Frankie Jackson, David Loope, and Aubrey Shirk (Chapter 11). The Cenozoic is well represented by six different trips. Nick Lang, B.J. Walker, Lily Claiborne, Calvin F. Miller, Rick Hazlett, and Matt Heizler (Chapter 9) examine spectacular cross-section view of the Miocene Spirit Mountain batholith and a coeval, and possibly related, eruptive center (Secret Pass) in the Colorado River extensional corridor. Another volcano-plutonic complex, the River Mountains–Wilson Ridge igneous system, which was dismembered by the Saddle Island detachment fault is the destination of a trip led by Denise Honn and Gene Smith (Chapter 1). Jim Faulds, Keith Howard, and Ernie Duebendorfer examine synextensional basins that constrain the timing of the structural demarcation between the Colorado Plateau and the Basin and Range (Chapter 6). Jerry Osborn, Matthew Lachniet, and Nick Saines weigh the evidence for and against Pleistocene glaciation in the Spring Mountains of southern Nevada in Chapter 7. The cultural effects of some of the youngest volcanism in the continental United States outside the Cascades is the focus of a trip by Sarah Hanson, Wendell Duffield, and Jeffrey Plescia (Chapter 8) to the San Francisco volcanic field near Flagstaff, Arizona. Finally, Kurt Frankel and a cast of thousands bring us up to date with a look at the active tectonics of the eastern California shear zone with discussions regarding significant discrepancies between long-term slip rates and the current rate of strain accumulation along active faults (Chapter 3). With field trips ranging from old to the present, the middle crust to the surface, from tectonics to paleontology, and from volcanism to glaciation, this volume offers something for everyone. Ernest M. Duebendorfer Eugene I. Smith

v

Map of the Nevada, California, Arizona, and Utah areas visited in these field trips showing locations of trips by number.

The Geological Society of America Field Guide 11 2008

The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section Denise Honn* Eugene I. Smith* Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154-4010, USA

ABSTRACT This ﬁeld trip will visit the River Mountains volcanic section (12.17 ± 0.02 to 13.45 ± 0.02 Ma) and Wilson Ridge pluton (13.10 ± 0.11 Ma) in southern Nevada and northwestern Arizona. Although this volcanic-plutonic system was disrupted by the Saddle Island detachment fault during Miocene crustal extension, there are convincing lithological, mineralogical, geochemical and geochronological indicators that suggest a cogenetic relationship. The trip consists of 17 stops that emphasize evidence that links the volcanic and plutonic sections. In addition we will visit the Saddle Island detachment fault at its type locality on Saddle Island. Keywords: plutonic rocks, volcanoes, Lake Mead, petrology, geochronology. The River Mountains volcanic section–Wilson Ridge pluton igneous system crops out at the northern end of the Colorado River extensional corridor, a north-south trending area of southern Nevada, western Arizona and eastern California that underwent up to 100% extension between ca. 23 and 12 Ma. In the northern part of the corridor, volcanic rocks of Tertiary age lie on Precambrian crystalline rocks and locally a thin conglomerate containing sedimentary and crystalline clasts. Paleozoic and Mesozoic sedimentary sections are missing and were probably stripped from a rising structural arch (the Kingman Arch) during late-Cretaceous, early Tertiary time (Faulds et al., 2001). The arch plunges gently to the north (~15°) and terminates against the Lake Mead fault system just north of Lake Mead. In the Colorado River extensional corridor, magmatism migrated to the north, pre-dating crustal extension by about 1 m.y. (Faulds et al., 2001).

INTRODUCTION The study of an igneous system is limited by exposure and preservation of the rock record. In most cases, only a portion of the system is exposed (i.e., volcanic or plutonic) and therefore only part of the magmatic history can be studied. Based on work done over the past 20 years, we interpret the River Mountains volcanic section of southern Nevada and the Wilson Ridge Pluton in northwestern Arizona as volcanic and plutonic segments of the same igneous system (Fig. 1). The connection between the River Mountains volcanic section and the Wilson Ridge pluton is based on structure, lithology, mineralogy, geochemistry, and geochronology. This ﬁeld trip will visit both the River Mountains and Wilson Ridge and will emphasize links between the volcanic and plutonic sections. *[email protected], [email protected]

Honn, D., and Smith, E.I., 2008, The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 1–20, doi: 10.1130/2008.ﬂd011(01). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

1

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Figure 1. Geologic map of Lake Mead region. Adapted from Smith et al. (1990).

Boulder City

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2

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section Quaternary

3

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Faults - dotted where concealed or inferred ball on downthrown side

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Contacts

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Tgp - Unassigned volcanic rocks

Tertiary

Tvu - upper Hoover Dam volcanic rocks Powerline road volcanic rocks Tpd - dacite Tpd - basalt and andesite Tbc - Boulder City Pluton Tbv - Bootleg Wash volcanic rocks Thd - Tuff of Hoover Dam and Dam Conglomerate Tbw - Boulder Wash volcanic rocks Tbr - Breccia River Mountains Stratovolcano Trs - quartz monzonite stock Tsv - andeste and dacite

Wilson Ridge Pluton Twrh - hypabyssal phase Twrm - medium grained phase Twrc - coarse grained phase Twrg - red feldspar granite phase Tid - diorite phase

Tpm - Patsy Mine volcanic rocks K-Tpp - Paint Pots intrusive rocks

Mesozoic Mz-p - Pennsylvanian through Mesozic rocks Paleozoic Precambrian

M-Pc - Precambrian through Mississippian rocks

Figure 1 (continued).

RIVER MOUNTAINS VOLCANIC SECTION The River Mountains volcanic section (12.17 ± 0.02 to 13.45 ± 0.02 Ma, 40Ar/39Ar whole-rock and mineral dates; Faulds et al., 1999) composed of mainly dacite, andesite, basalt and rhyolite is locally intruded by hypabyssal dacite plugs and a quartz monzonite stock. Smith (1982; 1984) suggested that the River Mountains are composed of at least four volcanoes that were juxtaposed by mid-Tertiary strike-slip faulting related to the leftlateral Lake Mead fault system: 1. The River Mountains stratovolcano and related satellitic dacite, rhyolite and basalt volcanoes. The stratovolcano is cored by the River Mountains quartz monzonite stock, which is surrounded by a zone of altered volcanic rocks cut by numerous dikes. The stock contains many xenoliths of basalt and dolomite. Dikes of porphyritic dacite radiate from the plug. The stock is chemically equivalent to rocks of the Wilson Ridge pluton and may represent the detached apex of one of the Wilson Ridge intrusions. Rocks above the intrusion are altered and mineralized andesite and plutonic rock cut by numerous dacite dikes that emanate from the River Mountains stock. The blue sodium amphibole, magnesio-riebeckite, occurs along

fractures and coatings on rocks of the quartz monzonite stock and surrounding altered volcanic rocks. Magnesioriebeckite is also found in fractures and thin veins in various phases of the Wilson Ridge pluton and within the Colorado River extensional corridor appears to be unique to this volcanic-plutonic system. 2. The Bootleg Wash section just north Boulder City, Nevada, composed from base to top of a section of andesite ﬂows, volcaniclastic breccia, and ﬂow-banded dacite ﬂows. 3. The Red Mountain section formed by highly altered andesite and dacite ﬂows, volcaniclastic rocks, and local granitic intrusions. On Red Mountain, andesite ﬂows and breccia are interleaved along numerous low-angle faults. The Red Mountain section is separated from the River Mountains stratovolcano by a northwest-striking fault (probably strike slip) and may represent highly altered volcanic and plutonic rocks related to the Boulder City pluton (14.17 ± 0.6 Ma; NAVDAT (http://navdat. kgs.ku.edu/); 13.8 Ma K/Ar age reported by Anderson et al., 1972). 4. The Casino dacite just east of Railroad Pass is characterized by andesite and dacite ﬂows and a thin moderately welded ash-ﬂow tuff.

4

Honn and Smith

WILSON RIDGE PLUTON (ABSTRACTED FROM LARSEN AND SMITH, 1990) The Wilson Ridge pluton is an epizonal to hypabyssal calcalkaline pluton that formed ca. 13.10 ± 0.11 Ma (40Ar/39Ar hornblende date; Faulds et al., 1999) during a period of mid-Miocene extension. Faulting and erosion have provided a cross section of the pluton in plan view. Geobarometric data and geologic constraints indicate the pluton has been tilted 17° to the north (Metcalf et al., 1993). The apex of the pluton, just south of Boulder Canyon (Lake Mead), Nevada, is composed of hypabyssal quartz monzonite and dacite cut by numerous dikes of rhyolite, dacite and basalt. The base of the pluton is 20 km to the south where quartz monzodiorite, monzodiorite, and diorite are in low-angle intrusive contact with Precambrian basement. The pluton was separated from comagmatic volcanic rocks in the River Mountains by movement along the Saddle Island fault system which includes the Saddle Island detachment, Hamblin Bay, and Eldorado faults (Weber and Smith, 1987; Duebendorfer et al., 1990). The age of detachment is estimated to be younger than ca. 13.5 Ma based on the inference that detachment faulting must postdate the formation of the Wilson Ridge pluton and River Mountain volcanic suite (Duebendorfer et al., 1990). The River Mountains now lie approximately 20 km to the west of the pluton. The Wilson Ridge pluton is composed of the Teakettle Pass suite consisting of foliated monzodiorite and quartz monzodiorite, unfoliated quartz monzonite, and the older Horsethief Canyon diorite. The Teakettle Pass suite comprises the main phase of the Wilson Ridge pluton (80 km2 outcrop area). The major minerals of the coarse-grained quartz monzonite (the dominant phase of the Teakettle Pass suite) are quartz (20%), orthoclase (25%), plagioclase (40%), and subhedral prismatic hornblende (<5%) (Larsen, 1989). The secondary phase of the Teakettle Pass suite is mediumgrained quartz monzodiorite and monzodiorite. Major minerals of this phase are plagioclase (45%), interstitial quartz and orthoclase (<20%), subhedral to euhedral biotite and hornblende (40%) (Larsen, 1989). Accessory minerals ubiquitous to the Teakettle Pass suite include sphene (2–4%), apatite, and zircon. The Teakettle Pass suite intrudes the Horsethief Canyon diorite in the southern portion of the Wilson Ridge pluton. The Horsethief Canyon diorite (4 km2 outcrop area) is composed of

plagioclase (50%), hornblende (35%), biotite (10%), anhedral quartz and orthoclase (<10%), megascopic 1–4 mm diameter sphene (2%–4%), and trace amounts of apatite (Larsen, 1989). The diorite is also present as angular to rounded xenoliths within the Teakettle Pass suite. Intermediate rocks of the Teakettle Pass suite contain abundant basalt and diorite enclaves. Basaltic enclaves are lensoidal and pillow-like and commonly have crenulate and ﬁne-grained margins. The enclaves probably represent blobs of maﬁc liquid that commingled and mechanically mixed with felsic magma to produce the intermediate rocks of the pluton. Basaltic enclaves commonly occur as inclusion-rich zones that represent synplutonic maﬁc dikes injected into a quartz monzonite host. Maﬁc magma was entrained and mechanically broken down by magmatic ﬂow shear. A continuum in shape exists from enclaves that are bulbous and ellipsoidal to thin, tabular maﬁc selvages and schlieren and ultimately to the maﬁc component in foliated quartz monzodiorite and monzodiorite (Larsen, 1989). Diorite enclaves have angular contacts with host rocks and are interpreted as xenoliths. Saddle Island Detachment Fault The Saddle Island detachment fault (Smith, 1982) cut the River Mountains volcanic section–Wilson Ridge pluton igneous system and moved the River Mountains volcanic section approximately 20 km to the west of the Wilson Ridge pluton (Weber and Smith, 1987) (Fig. 2). The detachment is exposed on Saddle Island on the west shore of Lake Mead just east of the River Mountains. Based on lithology and structure, Sewall (1988) divided the upper plate of the Saddle Island detachment fault into three fault-bounded domains. Domain one includes the Lower Cambrian Tapeats Sandstone and Pioche Shale and the overlying Miocene Horse Spring Formation. This lithology of this domain is very similar to exposures, interpreted as the upper plate of the Saddle Island detachment, on the east side of Wilson Ridge at the Cohenour Mine and in Petroglyph Wash (Feuerbach, 1986). On Saddle Island domain one is cut by mid-Miocene hypabyssal dacite and hornblende dacite. Domain two includes Precambrian amphibolite, gneiss, mica schist, quartz monzonite, and pegmatite intruded by mid-Miocene dacite, hypabyssal dacite, and diorite. The Saddle Island detachment is the lower boundary

W

E Projection of Saddle Island Detachment River Mountains

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Arch Petroglyph Mountain Wash

Figure 2. Cross section through from the River Mountains on the west to Detrital Wash on the east. The Saddle Island detachment fault crops out on Saddle Island east of the River Mountains and on Arch Mountain east. Adapted from Duebendorfer et al. (1990).

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section

Evidence for the Volcano-Pluton Link Based on lithology, mineralogy, geochemistry, geochronology, and structure, the following observations support a link between Wilson Ridge and the River Mountains. 1. Maﬁc enclaves are present in the Teakettle Pass suite and Horsethief Canyon diorite of the Wilson Ridge pluton. Larsen (1989) and Larsen and Smith (1990) demonstrated that these enclaves are chemically similar to maﬁc dikes of the Wilson Ridge pluton and to basalt ﬂows of the River Mountain volcanic section (Fig. 3) (Table 1). 2. Relatively immobile trace elements Th (4.2–25.1 ppm), Hf (3.3–6.6 ppm), and Ta (0.9–1.7 ppm) from 88 samples from the pluton and 32 samples from the volcanic section are tightly clustered and overlap (Fig. 4) (Table 1). Samples from nearby igneous systems plot outside of this cluster (Duebendorfer et al., 1990; Feuerbach, 1986; Larsen, 1989; Larsen and Smith, 1990). Also, chondrite-normalized rare-earth element distributions for Saddle Island dacite and Wilson Ridge volcanic and plutonic rocks overlap but are different from rareearth element values for the nearby Boulder City pluton (Duebendorfer et al., 1990) (Fig. 5). On a 87Sr/86Sr versus SiO2 plot, rocks of the Wilson Ridge pluton and the River Mountains volcanic suite form a linear trend with 87Sr/86Sr increasing with increasing SiO2 (Fig. 6) (Duebendorfer et al., 1998). 3. Rocks of the River Mountains and Wilson Ridge pluton have been dated using several techniques, including KAr (sanidine, biotite and whole rock), 40Ar/ 39Ar (sanidine, biotite, hornblende, and whole rock) and 206U/238Pb (zircon). 40Ar/ 39Ar dates from the Wilson Ridge pluton and River Mountains volcanic suite overlap and range from 13.45 ± 0.02 to 12.17 ± 0.02 (40Ar/ 39Ar whole-rock and mineral dates; Faulds et al., 1999). Preliminary 206Pb/238U zircon dates (LA-ICPMS) from 106 to 40 µm spots on 49 zircons suggest a complex multiphase system active

Sun/McDon. 1989-PM 100,000

10,000 Rock/Primitive Mantle

of this domain. Domain three includes Precambrian basement similar to that of domain two; however, domain three has few mid-Miocene intrusions. Dacite and hypabyssal dacite intrusions of Saddle Island contain rounded porphyritic basaltic enclaves commonly with embayed and cuspate margins. This texture resulting from magma commingling is common within both the River Mountains volcanic section and in the southern part of the Wilson Ridge pluton. The zone of detachment faulting between the upper and lower plates consists of a 2-m-thick microbreccia or ultracataclasite that overlies a 30-m-thick zone of chlorite phyllonite. The lower plate contains variably mylonitized amphibolite, gneiss, and pegmatite. The degree of shearing (mylonitization) in the lower plate increases upward toward the detachment zone. A complete description of the Saddle Island detachment is provided in Duebendorfer et al. (1990).

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basalt from the RM Mafic dikes of the WRP basaltic enclaves WRP

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WRP RM BCP

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Ta

Figure 4. Ternary diagram reﬂecting the tight clustering of immobile elements Th, Hf, and Ta from the Wilson Ridge pluton (WRP) and River Mountains volcanic section (RM). The Boulder City pluton (BCP) does not ﬁt in this tight cluster. Data from Duebendorfer et al. (1990); Feuerbach (1986); Smith et al. (1990); and Weber and Smith (1987).

for 4.2 million years (based on a zircon core-rim pair) to a maximum of 7.2 million years (based on two zircon rim dates 18.9 ± 0.8 to 13.1 ± 0.6 Ma) (Honn et al., 2007) (Fig. 7). K-Ar and 40Ar/ 39Ar sanidine, biotite, and hornblende whole-rock dates represent emplacement ages within the crust and therefore reﬂect a short history for the system. In contrast, 206U/238Pb zircon dates reﬂect the entire history as well as the age of antecrysts and

66.7 90.3 21.2 4.4 0.8 0.5 1.4 0.3 17.2 16.8 5.4 1989 NA NA 1.4 3.7

Trace elements in ppm La 88.9 97.9 Ce 105.7 145.8 Nd 41.0 61.0 Sm 6.5 8.6 Eu 1.4 1.6 Tb 0.5 1.1 Yb 1.5 2.0 Lu 0.4 0.4 U 16.9 15.2 Th 20.1 18.9 Hl 5.9 8.1 Sa 1840 1146 Rb 77 151 Sr 410 680 Ta 1.7 1.9 Sc 4.0 5.7

0.4

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101.7

0.2

GWR WR46 68.5 14.0 4.4 2.2 1.1 4.1 5.5 0.6 0.1 0.3

100.8

LOI

101.0

0.1

Oxide Si02 AI203 Fe203 CaO MgO Na20 K20 Ti02 MnO P205

Total

GWR WR368 68.2 15.4 4.6 1.5 1.0 5.2 4.5 0.4 0.0 0.1

46.6 59.8 26.2 4.7 1.0 0.4 1.6 0.3 11.1 11.7 4.8 1274 112 249 1.3 5.4

101.0

0.1

GWR WR55 68.1 14.4 6.0 2.0 1.0 4.6 4.3 0.3 0.0 0.2

64.2 88.6 24.4 4.1 0.6 0.4 1.4 0.3 18.8 21.0 4.9 811 129 NA 2.2 2.3

100.9

0.1

GWR WR57 74.3 13.2 2.6 0.9 0.4 4.4 4.7 0.2 0.0 0.1

80.9 117.0 34.2 6.5 1.3 0.5 1.5 0.3 11.9 15.4 5.5 1796 89 314 1.4 4.5

100.3

0.0

48.2 57.6 19.0 3.4 0.7 0.9 1.1 0.2 10.3 16.3 3.6 1021 108 163 1.5 2.7

99.9

0.0

49.7 59.3 18.3 4.0 0.8 0.8 0.9 0.3 11.3 17.2 5.3 1702 78 NA 1.5 3.6

102.0

0.7

Wilson Ridge pluton GWR GWR GWR WR58 WR62 WR63 67.0 70.2 71.3 15.2 13.8 14.2 5.3 4.7 4.2 1.8 1.1 0.8 0.9 0.5 1.0 5.4 4.7 5.5 4.1 4.6 3.9 0.4 0.2 0.3 0.0 0.0 0.0 0.2 0.1 0.1

62.6 89.8 25.1 4.4 0.7 0.3 1.1 0.2 20.5 19.1 4.1 975 164 NA 1.3 1.5

102.1

0.3

GWR WR67 72.7 13.1 4.6 0.1 0.2 0.6 10.4 0.2 0.0 0.0

64.2 88.6 24.4 4.1 0.6 0.4 1.4 0.3 18.8 21.0 4.9 811 129 NA 2.2 2.3

102.1

0.2

GWR WR56 65.6 15.7 5.5 3.2 1.5 4.7 4.7 0.6 0.0 0.4

58.3 98.6 37.9 5.6 1.1 0.5 1.9 0.3 1.6 16.0 4.9 1008 125 NA 1.5 4.1

102.1

1.1

GWR WR71 71.7 14.4 3.1 0.8 0.3 3.9 6.4 0.3 0.0 0.1

17.4 26.7 10.5 2.1 0.4 0.3 1.7 0.3 2.4 17.2 5.5 903 96 99 1.2 2.2

100.3

0.2

GWR WR65 71.3 14.2 3.6 0.5 0.2 5.3 4.8 0.2 0.0 0.0

101.2

0.1

GWR WR36 71.0 14.2 4.5 0.7 0.6 4.4 5.4 0.3 0.0 0.1

54.0 55.8 89.8 95.4 18.7 8.1 4.5 5.0 1.0 1.0 0.4 0.4 1.5 1.8 0.3 0.2 0.8 2.4 12.2 16.9 4.3 4.9 736 1554 54 143 264 230 1.1 1.4 3.9 3.6 (continued)

101.1

0.7

GWR WR72 71.6 14.9 3.5 0.7 0.7 7.3 1.5 0.3 0.0 0.0

TABLE 1. CHEMICAL ANALYSES OF REPRESENTATIVE IGNEOUS ROCKS FROM THE WILSON RIDGE PLUTON AND THE RIVER MOUNTAINS

6 Honn and Smith

47.9 77.1 20.5 3.8 0.8 0.2 1.3 0.2 1.9 13.5 4.3 1281 121 254 1.1 2.6

Trace elements in ppm La 49.7 51.9 Ce 79.0 89.5 Nd 31.9 35.2 Sm 5.9 5.1 Eu 1.4 1.0 Tb 0.5 0.5 Yb 1.8 2.1 Lu 0.3 0.3 U 1.4 2.2 Th 10.0 13.7 Hl 5.1 4.6 Sa 1448 1063 Rb 93 217 Sr 646 212 Ta 1.3 1.2 Sc 7.1 4.2

0.3

GWR WR33 73.7 14.2 2.7 1.1 0.6 4.8 4.5 0.2 0.0 0.0

102.0

0.6

GWR WR67 71.5 13.4 3.8 0.4 0.6 1.1 9.7 0.2 0.0 0.1

101.5

101.8

0.0

LOI

Total

GWR WR57 67.6 15.6 5.2 3.0 1.2 4.8 3.7 0.4 0.0 0.2

Oxide Si02 AI203 Fe203 CaO MgO Na20 K20 Ti02 MnO P205

28.8 43.6 12.6 1.6 0.2 0.1 1.2 0.3 4.2 25.1 3.3 123 164 NA 1.7 1.1

100.7

0.2

GWR WR59 77.5 12.1 2.1 0.4 0.1 4.3 4.0 0.1 0.0 0.0

TABLE 1. (continued)

28.8 43.6 12.6 1.6 0.2 0.1 1.2 0.3 4.2 25.1 3.3 123 164 NA 1.7 1.1

100.7

0.2

10.3 143.2 43.4 0.8 1.6 0.5 2.4 0.0 0.2 18.9 7.7 166 64 339 1.4 4.4

101.2

0.4

69.8 99.2 24.5 4.5 0.9 0.4 1.6 0.3 2.1 18.9 5.2 1045 113 227 1.3 2.8

101.2

0.2

58.6 91.2 38.1 6.1 1.4 0.7 1.7 0.3 1.0 8.8 5.0 1679 87 524 1.0 6.4

99.5

0.2

Wilson Ridge pluton (continued) GWR GWR GWR GWR WR59 WR106 WRl14 WR120 77.5 68.0 72.2 65.1 12.1 15.8 14.3 16.5 2.1 4.4 2.5 4.0 0.4 1.6 1.0 3.2 0.1 1.0 0.5 1.5 4.3 5.4 3.6 4.0 4.0 4.1 6.6 4.4 0.1 0.4 0.2 0.5 0.0 0.0 0.0 0.1 0.0 0.0 0.0 0.2

206.5 328.1 101.0 13.6 3.5 1.3 2.2 0.3 NA 17.0 10.4 2830 63 1660 0.9 8.7

101.4

0.4

DWR WR74 59.7 19.1 6.2 4.7 1.7 6.1 2.6 0.7 0.1 0.2

90.5 173.2 68.2 12.5 2.4 1.4 3.1 0.4 2.4 14.4 7.8 1339 126 415 1.8 17.9

99.4

1.8

DWR WR57d 53.6 15.4 8.9 5.5 4.5 4.6 3.0 1.1 0.2 0.8

50.3 79.0 39.0 7.3 2.0 1.2 1.6 0.2 NA 4.5 4.0 1055 77 778 0.6 22.7

101.8

1.5

DWR WR39d 57.0 15.6 6.6 6.7 6.5 4.7 2.2 0.7 0.1 0.3

105.0 167.1 52.7 11.2 2.6 1.3 2.0 0.3 2.0 11.0 6.0 1883 70 663 0.9 11.1

99.0

1.6

DWR WR107 55.9 17.4 6.4 4.2 2.8 5.2 4.2 0.8 0.1 0.5

54.0 76.4 NA 8.4 2.1 1.1 2.0 0.4 1.0 7.0 5.0 1632 54 505 1.3 22.7 (continued)

98.0

0.4

GWR WR121 53.3 17.6 7.6 6.7 3.5 3.7 3.3 1.4 0.1 0.4

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section 7

2.3

99.2

Oxide Si02 AI203 Fe203 CaO MgO Na20 K20 Ti02 MnO P205

LOI

Total

99.8

1.7

PB 219 46.0 14.3 12.9 8.6 8.4 3.3 1.5 2.1 0.2 0.7

100.0

0.3

DP 220 67.1 14.4 5.0 3.1 0.9 4.7 3.8 0.5 0.1 0.2

100.0

3.0

DP 223 67.2 14.2 3.5 2.3 0.9 3.6 4.9 0.4 0.1 0.1

100.4

1.0

DP 271 66.0 14.5 4.6 3.8 1.0 3.5 5.2 0.5 0.1 0.2

99.0

0.1

RP 348 72.9 12.6 2.7 0.5 0.2 2.2 7.5 0.2 0.0 0.0

100.8

4.4 98.8

0.2

River Mountains RP AP 349 346 71.2 59.1 13.3 16.5 1.8 6.4 1.3 4.9 0.5 1.7 3.0 4.2 4.8 3.9 0.2 0.9 0.1 0.1 0.0 0.6

99.8

1.7

AP 347 61.7 14.5 6.0 3.9 1.2 2.6 6.9 0.7 0.1 0.3

99.4

1.0

AP 222 59.0 16.6 6.9 5.5 2.6 3.9 2.3 0.9 0.1 0.6

99.6

0.8

DB 229 67.7 12.2 5.6 2.1 0.4 2.4 7.7 0.6 0.0 0.2

98.9

3.3

AB 230 59.5 13.1 6.9 4.6 1.4 3.0 5.3 1.1 0.1 0.5

99.1

4.7

AB 226 53.5 15.0 7.8 5.2 3.8 4.1 3.2 1.1 0.1 0.7

99.4

0.8

MR 231 64.9 15.2 5.3 1.0 1.7 4.7 5.0 0.6 0.1 0.2

99.5

1.3

DR 229 65.2 14.2 5.2 1.6 1.7 2.5 7.2 0.5 0.1 0.1

98.8

1.4

DR 272 56.6 17.0 5.8 4.4 3.0 4.0 5.2 0.8 0.1 0.5

Trace elements in ppm La 107.6 75.9 64.7 52.4 55.0 55.6 59.5 93.5 82.1 81.7 47.7 53.6 74.5 63.0 49.9 75.9 Ce 242.4 170.9 119.3 90.7 105.4 124.5 124.5 203.5 181.5 167.9 102.6 120.3 163.6 121.6 97.6 157.2 Nd 117.7 44.2 41.3 34.2 37.0 22.4 29.1 50.6 37.6 63.3 38.5 56.5 75.5 49.3 35.6 77.9 Sm 15.5 9.1 5.6 4.2 5.6 3.7 4.6 7.5 7.2 8.7 5.8 7.5 10.6 5.6 4.5 9.1 Eu 3.9 2.7 1.4 1.2 1.3 1.0 1.0 1.9 1.9 2.0 1.1 2.1 2.5 1.4 1.1 2.2 Tb 3.0 2.0 0.9 0.3 1.2 0.4 0.4 0.9 0.4 1.5 0.3 0.5 0.7 0.1 0.2 1.1 Yb 2.1 1.7 1.5 1.3 1.6 2.2 2.2 2.5 2.2 2.3 1.4 2.0 2.4 1.5 1.4 1.7 Lu 0.4 0.3 0.3 0.2 0.3 0.3 0.3 0.4 0.3 0.4 0.2 0.3 0.4 0.3 0.3 0.3 U 6.4 8.7 9.9 9.1 7.6 3.9 3.6 2.1 2.0 4.6 3.0 2.6 2.0 NA 1.3 1.3 Th 13.2 7.7 18.5 17.6 15.4 31.2 26.3 19.3 18.6 14.9 14.8 10.8 20.7 14.2 15.8 15.7 Hl 8.1 6.4 4.7 4.8 4.7 4.9 4.5 9.2 6.7 7.3 4.1 5.1 6.6 5.5 4.5 5.8 Ba 2155 891 1484 1470 1517 650 1510 2740 1930 1650 774 1121 1376 1950 1683 1625 Rb 58 33 152 187 163 166 166 155 205 121 199 161 99 118 184 139 Sr 1205 614 230 131 175 110 200 470 420 504 30 33 351 33 60 405 Ta 1.5 1.6 0.9 1.2 1.1 15.3 14.6 15.1 11.1 1.0 1.2 1.1 0.9 1.0 0.9 1.1 Sc 29.6 27.4 6.5 3.8 6.9 1.5 1.5 5.2 8.7 11.2 7.5 14.9 14.6 6.9 5.9 11.0 Note: Major elements in weight percent. NA—not analyzed or not detected. Total iron as Fe203'. THD—Tuff of Hoover Dam AHD—Hoover Dam andesite. GWR—Wilson Ridge granite or quartz monzonite DHD—Hoover Dam dacite. DWR—Wilson Ridge diorite AP—Powerline Road andesite. PB—Powerline Road basalt. DB—Bootleg Wash dacite. DP— Powerline Road dacite. AB—Bootleg Wash andesite. MR—Quartz monzonite of the River Mountains stock. DR—Dacite of the River Mountains stratovolcano. AR—Andesite of the River Mountains stratovolcano.

PB 218 47.1 14.5 10.9 11.2 6.3 2.8 1.2 1.8 0.1 1.0

TABLE 1. (continued)

8 Honn and Smith

1000

REE / Chondrite

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section SiO 2 Saddle Island dacite 57.46% 66.96% 100

Boulder City Pluton

Figure 5. Rare earth element distribution diagram for the Wilson Ridge pluton (light stippling), Boulder City pluton (heavy stippling) and hypabyssal dacites from Saddle Island (line). SiO2 ranges are given for Boulder City pluton and Wilson Ridge pluton samples. Adapted from Duebendorfer et al. (1990).

SiO2 53.60% 77.48%

10

9

Wilson Ridge Pluton 1 La Ce

Nd

Sm Eu

Tb

Yb Lu

-6

0.7105 River Mtns. Wilson Ridge White Hills Aztec Wash Boulder City

-7

0.7095 (87Sr / 86Sr)

epsilon Nd

-8

River Mountains Wilson Ridge

0.7100

-9 -10 -11

0.7090 0.7085 0.7080 0.7075

-12

(A.)

-13 0.706

0.7070

(B.)

0.7065 0.707

0.708

0.709

0.710

0.711

45

0.712

50

(87Sr/86Sr)

55

60

65

70

75

SiO2

Figure 6. (A) The River Mountains volcanics and Wilson Ridge pluton rock are distinct from nearby volcanic and plutonic rocks in terms of their 87Sr/86Sr and epsilon Nd. (B) The positive correlation between SiO2 and 87Sr/86Sr suggests that the River Mountains volcanics and the Wilson Ridge pluton are cogenetic. This single linear array suggests magma mixing between two endmembers created the volcanic-plutonic suite. Adapted from Duebendorfer et al. (1998).

RM and WRP 40Ar/39Ar ages WRP late stage dike WRP medium grained phase WRP coarse grained phase Horsethief Canyon Diorite RM volcanic rock

(Millions of Years Ago)

10

15

20

25

30

35

40

45

RM stock

Figure 7. Preliminary LA-ICPMS U-Pb dates from zircon for the Wilson Ridge pluton (WRP) and River Mountains (RM). Dates are based on individual 40 micron spots. Dates older than 19 Ma are considered xenocrysts.

10

Honn and Smith

A

B Figure 8. Cathodoluminescent images of zircons used in U-Pb LA-ICPMS dating with xenocrystic cores. (A) Zircon has at least two major dissolution events (dashed lines). (B) Zircon has an angular core (dashed line) and small (10 microns in diameter) melt inclusion (black oval).

100 microns

100 microns

A F E

Hypabyssal and Volcanic Rocks

D

Quartz Monzonite

B

C Small Diorite Pluton

Zone of Hybridization

Ponded Basalt

Figure 9. Cartoon of petrogenetic model for the evolution of the River Mountains–Wilson Ridge igneous system (adapted from Larsen, 1989; Larsen and Smith 1990) with general stop areas within the system (A) River Mountains overlook area, (B) Kingman Wash, (C) Horsethief Canyon, (D) River Mountains stock and border zone, (E) Saddle Island detachment fault, (F) northern River Mountains volcanic section.

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section xenocrysts. Textural evidence in cathodoluminescence images was used to recognize xenocrysts and inherited cores (Fig. 8), which give ages from 1517.5 ± 11.2 Ma to 21.3 ± 0.8 Ma. Many of these older ages have larger error bars and likely reﬂect analytical errors (Honn et al., 2007). 4. The rare occurrence of magnesio-riebeckite (a sodic amphibole) in the Wilson Ridge pluton was studied by Potts (2000), but has only recently been noted in the River Mountains volcanic section. In both areas, magnesioriebeckite occurs as coatings on fracture surfaces, veinlets, dense stockwork veining, and diffusive alteration fronts. Potts (2000) concluded that magnesio-riebeckite in the Wilson Ridge pluton formed due to subsolidus Na-metasomatism. The metasomatism was related to hydrothermal alteration of the pluton involving Na-rich ﬂuids with meteoric origins (based on δ18O values of the amphibole). The timing of magnesio-riebeckite formation in the pluton is constrained by the high-angle faults that were active about 10 Ma, the surfaces of which are coated with magnesio-riebeckite, and the emplacement of the Fortiﬁcation Hill Basalt (4.61–5.89 Ma), which contains no magnesio-riebeckite (Potts, 2000). CONCLUSIONS Linked by structure, rock type, enclaves, geochemistry, and geochronology, the River Mountains volcanic suite and Wilson Ridge pluton are a rare example of a well-preserved and exposed igneous system that can be used to study magma chamber processes. Figure 9 is a schematic diagram from Larsen (1989) of the combined River Mountains–Wilson Ridge pluton igneous system that has been modiﬁed to show the position of ﬁeld trip stops within the River Mountains volcanic section–Wilson Ridge pluton magmatic system.

11

From the peak, the eastern skyline shows, from north to south, the River Mountains stratovolcano and the River Mountains stock. Across Lake Mead is the Wilson Ridge pluton. The prominent ﬂat-topped mesa in front of the Wilson Ridge pluton is Fortiﬁcation Hill, capped by basalt dated at 5.89 ± 0.08 Ma (Feuerbach et al. 1991). Tomorrow morning, we will be driving in Kingman Wash along the south side of Fortiﬁcation Hill. Andesite and dacite ﬂows form the ﬂanks of the River Mountains stratovolcano and are cut by dacite sills and dikes that radiate from the River Mountains stock. Many of the dikes grade from ﬁne-grained quartz monzonite near the stock to dacite at their distal ends. Alteration of the stratovolcano is primarily argillic and ferric; mineralization is characterized by barite, specular hematite, and manganese oxide. Although the Paleozoic section is generally missing (perhaps eroded during the uplift of the Kingman Arch), remnants of a Paleozoic carbonate section of Cambrian and Mississippian age is preserved and intruded by the quartz monzonite stock. Epidote, tremolite, and garnet are locally produced in these contact zones. Stop 2. Magnesio-Riebeckite exposures On our way back we will stop at an outcrop of magnesioriebeckite (sodic amphibole) within the River Mountains stratovolcano (Fig. 10). The Wilson Ridge pluton is the only other documented occurrence of magnesio-riebeckite in the Colorado River extensional corridor. In the Wilson Ridge pluton, magnesio-riebeckite formed by Na-metasomatism related to subsolidus hydrothermal alteration. δ18O (Potts, 2000) data suggest that the ﬂuids involved in metasomatism were meteoric (δ18O ranging from −5 to +10). Tomorrow we will see the extensive riebeckite alteration in the Wilson Ridge pluton. DAY 2 Area B. Kingman Wash

DAY 1 Area A. River Mountains Trail Overlook This is a short (2.4 km) hike from the River Mountains Trail parking lot in Boulder City to a peak of in the southern River Mountains for an introduction to the local geology, the River Mountains stratovolcano, and an outcrop of magnesio-riebeckite alteration (Figs. 10 and 11). Stop 1. River Mountains Overlook Once the trail enters the canyon, it follows the contact between the Red Mountain volcanic section (west side) and the River Mountains stratovolcano on the right (east) side of the wash. The contact is a northwest-striking fault (probably strike slip). The Red Mountain volcanic section may represent highly altered volcanic and plutonic rocks related to the Boulder City pluton (13.88 ± 0.1 Ma; Faulds et al., 1999).

The purpose of this stop is to see the main phase of the pluton, a late stage dike, riebeckite alteration, and younger volcanic units. Starting at the Kingman Wash road located 4 km east of Hoover Dam. Turn off of U.S. 93 onto Bureau of Land Management (BLM)–approved backcountry road 70 and turn right onto road 70C. See Figures 10 and 12 for area and stop locations. Stop 3. Magnesio-Riebeckite Outcrops along the last mile (1.6 km) of road 70C (6.75 miles or 10.9 km from the turn off of U.S. 93) have extensive magnesio-riebeckite alteration on fracture surfaces and in veins. These exposures of magnesio-riebeckite are evidence for the subsolidus sodic alteration. Quartz monzonite of the pluton is ﬁner grained here than in the main phase of the pluton. Based on the amount of tilting of the pluton, we are in the mid-level of the pluton at this location but closer to the margins than in Stop 4 (Fig. 10).

Tpd

Tp

d

QTs

Tbv

QTs

Tsv

ATrs D

Tpd

Tsv

Tsv

Tpd

Qts

114 50'

o

Boulder City

Tpb

F

QTs

m

Tp

QTs

E

M-PC

Tgp

Tgp

0

0

Tgp

Tcf

Tcf

Lake Mead

1 Mi

1 Km

Tgp

Tcf

Tcf

Tgp

Tbc

Tpm

QTs

114 45'

o

Thd

N

Tcf

Tvu

Thd

K-Tpp

Tmf

QTs

Tgp

QTs

Tvu

Promontory Point

Tgp

QTs

Tgp

QTs

Mz-P

QTs

Tmf

Tpm

K-Tpp

Tvu

QTs Tvu

Thc

QTs

QTs

114 40'

o

QTs

C

Twrm

M-PC

QTs

B

Twrm

Tgp

Tid

Tmf

Tgp

o

Tid

Twrc

Tgp

Tmf

Qts

114 35'

Twrc

Tmf

e r t on Gilb any Tmf C Twrh

Twrh

Tgp

Thc

Tbr M-PC

Twrh Tgp

Twrh

Tgp

Twrh

s QT

Twrh

Thc

QTs

Tbr

-P Mz

Twrc

Twrc

Twrh

QTs

Thc

114o 35'

Twrc

QTs

M-PC

Tmf

Tmf

Tmf

Tmf

QTs

QTs

QTs

Lake Mead

Boulder Wash

QTs

Tbw

Thc QTs

Qts

36o 00'

36o 05'

36o 10'

Figure 10. Stop area locations on a generalized geologic map of the Lake Mead region. Adapted from Smith et al. (1990). See Figure 1 for color version of the same geologic map without area locations.

114 50'

o

d

Tpd

Tpd

Tpb

Tpd

Tpb

Tpd

Tpd

Tpd

Tpb

Tpd

Tp

QTs

QTs

Tpd

QTs

Tpb QTs Tpd

Tgp

Tgp

Tpb

35o 00'

36o 05'

Tpb

114o 50'

36o 10'

Tgp

Tgp

Tcf

Fortification Ridge Tid

Tcf

114o 40'

WilsonRidge

QTs

Saddle Island

M-PC

Twrc

M-PC

114o 45'

Hoover Dam

Twrc Tmf

Arch Mountain

114o 50'

Detrital Wash

12 Honn and Smith

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section

114o51.000' W

13

114o50.000' W

7 RM stock

8 border zone

Area D 36o00.000' N

Area A RM overlook

1

2

magnesio-riebeckite

93

N

0

MN 12.5

O

.5 0

.5

1 mi 1 km

Figure 11. Topographic map with locations of stops for areas (A) River Mountains (RM) trail overlook and (D) River Mountains stock and border zone. Stops 1, 2, 7, and 8: 1—River Mountains overlook; 2—magnesio-riebeckite outcrop; 7—River Mountains stock; 8—border zone.

70c

5

3

N

12.5

MN O

0

end of the road campsite

114o40.000' W

magnesio-riebeckite

view of Fortification Hill

114o41.000' W

0

4

.5

.5

114o39.000' W

1 km

1 mi

Figure 12. Topographic map with locations of stops for area (B) Kingman Wash. Stops 3, 4, and 5: 3—magnesio-riebeckite; 4—end of the road campsite; 5—view of Fortiﬁcation Hill.

70

bathroom

Area B

114o42.000' W

36o03.000' N 36o02.000' N

14 Honn and Smith

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section Stop 4. End of the Road Campsite At the end of road 70C, the main phase of the Wilson Ridge pluton is medium- to coarse-grained quartz monzodiorite. This is the mid-level of the pluton (Fig. 10); magma commingling textures are present but not as common as near the base of the pluton in Horsethief Canyon (Area C). A late-stage dike crops out approximately 100 m up the trail (the now closed road) to the north. The dike is porphyritic rhyolite with plagioclase phenocrysts up to 1 cm in length.

15

Stop 5. View of Fortiﬁcation Hill (Abstracted from Metcalf et al., 1993) Backtrack along road 70C. The black mesa to the northwest is Fortiﬁcation Hill. Pliocene basalts in the Lake Mead area deﬁne the Fortiﬁcation Hill volcanic ﬁeld (Feuerbach et al., 1993). Basalts of this ﬁeld are subalkalic and alkalic and erupted between 5.89 and 4.7 Ma (K-Ar plagioclase separate dates; Feuerbach et al., 1991). Magmatic activity in the ﬁeld ended with the formation of low-volume alkali basalt centers in northwest Arizona along U.S. 93 and in Petroglyph Wash and near Boulder Beach. The Fortiﬁcation Hill–Lava Cascade group forms a chain of at least 6 vents with a length of 25 km. The chain appears to be controlled by north-northwest–striking faults that bound the west side of the Black Mountains. Vents occur both on the margin and in the interior of the range. The escarpment that forms the cap of Fortiﬁcation Hill is composed of over 100 ﬂows of olivine basalt and scoria. Cinder cones aligned in a north-south direction are intruded by dikes and plugs and served as the source for the ﬂows. The Fortiﬁcation Hill cinder cones represent a node of intense volcanic activity along a north-trending en echelon dike system. Dike orientation is coplanar with the east-dipping midMiocene Fortiﬁcation fault zone.

The complete spectrum from discrete xenoliths to commingled magmas to completely mixed rock in the Teakettle Pass suite is exposed in this canyon. Enclaves are lensoidal, fusiformal, tabular, pillow shaped, commonly have crenulate margins and ﬁne-grained borders, and are locally boudinaged. Many enclaves show a weak internal foliation that is subparallel with the foliation in the host and with the mesoscopic foliation deﬁned by the alignment of the enclaves themselves. Enclaves are typically 20–50 cm long and are rarely isolated; more typically they cluster in enclaverich tabular zones that display strong ﬂow foliation near their margins. Enclave zones occur throughout the pluton in the intermediate phases of the Teakettle Pass suite. In two dimensions, enclave zones range from <1 m by 5 m to 10 m by 500 m. A continuum in shape exists from enclaves that are bulbous and ellipsoidal to those that are thin, tabular maﬁc schlieren and ultimately to the maﬁc component in foliated quartz monzodiorite and monzodiorite. Chilled borders, crenulate margins, boudinage, and pillowlike geometries of maﬁc enclaves of the Teakettle Pass suite strongly suggest that they were liquid or semi-liquid at the time of their incorporation in the felsic host. The abundance of lensoidal and phacoidal inclusions and the lack of strain within the inclusions and most phases of the Teakettle Pass suite indicate that enclaves were deformed while still partially molten. The lack of sharply bounded, tabular dikes with chilled margins suggests that dike injection occurred before the felsic host was crystallized enough to sustain a fracture. Dike emplacement occurred continuously as the pluton cooled. Late-stage dikes are tabular, cut all phases of the pluton except felsic dikes, and are colinear with high-angle normal faulting.

Area C. Horsethief Canyon

DAY 3

The purpose of this stop is to see the commingling textures near the ﬂoor of the pluton between the Teakettle Pass suite and Horsethief Canyon diorite. After returning to U.S. 93 drive 2 miles (3.2 km) east and turn onto on BLM-approved backcountry road make a left turn onto road 66 (just before the White Rock Canyon turn off) and continue to its end (4.75 miles; 7.64 km) at the dry falls. See Figures 10 and 13 for stop locations.

Area D. The River Mountains Stock and Border Zone

Stop 6. Campground at Dry Falls/End of the Road (Abstracted from Larson, 1989, Larson and Smith, 1990, and Metcalf et al., 1993) The spectacular magma-commingling textures on the dry falls at the entrance to the canyon are just a taste of what’s to come. We will use the trail up the hill on the north side of the dry falls to get into the canyon. We are now near the ﬂoor of the pluton (Figure 10). On a 1–2 km hike up the canyon we will see exposures of the Horsethief Canyon diorite intruded by basalt (now enclaves), magmas of the Teakettle Pass suite (quartz monzonite) as well as xenoliths of diorite within monzodiorite.

This is a short (1.6 km) hike to outcrops of the River Mountains stock, border zone, and an overlook of the River Mountains stratovolcano (Fig. 10 and 11). Stop 7. The River Mountains Stock (Abstracted from Smith, 1984) Heading east on U.S. 93 toward Hoover Dam from Boulder City, turn north (left) onto BLM-approved backcountry road 76 then turn left onto road 77 and continue on to the break in slope. The River Mountains stock is a plug in the conduit of the River Mountains stratovolcano. The stock is a composite pluton composed of ﬁne- to medium-grained quartz monzonite. Near the top of the stock the rock resembles ﬁne-grained dacite. The stock has radiating dikes of monzonite that intrude a border zone of altered and mineralized volcanic rock. The stock contains several large blocks (10 m in size) of Paleozoic carbonate, Tertiary andesite, and basalt.

93

Area C

N

12.5

MN O

6

0 0

.5

.5

campground at dry falls

114o39.000' W

Figure 13. Topographic map with locations of stop for area (C) Horsethief Canyon. Stop 6: campground at dryfalls.

66

114o40.000' W

1 km

1 mi

114o38.000' W

36o00.000' N 35o59.000' N

114o41.000' W

16 Honn and Smith

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section Stop 8. The Border Zone (Abstracted from Smith, 1984) Head back down road BLM-approved backcountry road 77 until the canyon narrows about (0.5 miles; 0.8 km). The border zone surrounding the River Mountains stock is a complex transition zone of several porphyritic dacite dikes that intrude the highly altered and mineralized andesite of the River Mountains stratovolcano as well as Paleozoic carbonate blocks. The contact between border zone and volcanic rocks of the River Mountains stratovolcano is gradational. Optional Area E. Saddle Island Detachment The purpose of visiting this area is to see outcrops of the Saddle Island detachment fault and rocks in the upper plate (Fig. 10 and 14). Optional Stop 9. Saddle Island Detachment Turn east off of Lake Shore Road onto the paved road to Southern Nevada Water System Treatment and Pumping Plant operated by the Southern Nevada Water Authority (SNWA). Not only does Saddle Island house the water treatment plant, two of the intakes for Las Vegas’ water supply are located on the east side of the island. Because of the strategic importance of the island, SNWA restricts access. We will visit Saddle Island only if we obtain permission from SNWA. After passing through the security checkpoint, drive across the causeway to Saddle Island and park at the turnout just before the road turns south. Optional Stop 11. Detachment Fault The purpose of this stop is to view the zone of intense shearing and the microbreccia just below the detachment fault. We will hike about 0.5 km to the northeast to the saddle at the crest of the Saddle Island ridge. The hike begins in weakly foliated Precambrian amphibolite cut by scattered dikes of muscovitebearing pegmatite. In some areas the Precambrian amphibolite is weakly foliated and contains hornblende, plagioclase, and chlorite as dominant minerals. In other areas, amphibolite displays a strong near-horizontal mylonitic foliation. Thin zones of brittle shearing are superimposed on the mylonitic fabric. In these zones, the rock develops a schistose fabric and earlier textures are rarely preserved. The reddish rocks on the north side of the saddle are brecciated Tertiary dacites that form the lowermost part of the upper plate on Saddle Island. The detachment fault is located about half way up the ridge at the color change (green rock = lower plate; red rock = upper plate). On the south side of the saddle, amphibolite displays a weak mylonitic foliation, but little evidence of brittle shearing is present. About 20 m below the fault, amphibolite begins to show the effects of brittle shearing. As the fault is approached, the lower-plate rock progressively becomes more intensely sheared and is converted into chlorite schist. The strike of the brittle foliation is N30E (coplanar with the detachment). Notice the small-scale, lens-like pattern of foliation and the boudin of quartz-feldspar rock (intensely sheared pegmatite dikes?).

17

Just below the fault, a black, ﬁne-grained microbreccia or ultracataclasite is exposed. This rock displays strong ferric alteration (hematite). The microbreccia marks the uppermost part of the lower plate and represents the most intense shearing along the detachment fault. The fault itself is not exposed in the saddle. There is an excellent exposure, however, on the east side of island, just below the saddle. Optional Stop 12. Upper Plate Rocks Continue walking to the north along the crest of the Saddle Island ridge. The three fault bounded domains of Sewall (1988) in the upper plate are divided into four lithological units (Duebendorfer et al., 1990). These are: (1) Precambrian amphibolite, schist, gneiss, and granite; (2) Precambrian basement intruded by dikes and small plugs of quartz monzonite and hypabyssal dacite and diorite; (3) a Lower Cambrian section consisting of the Tapeats Sandstone, Bright Angel Shale, and Bonanza King Dolomite that strike northwest and is vertical to overturned; and (4) conglomerate and megabreccia of Tertiary (?) age (possibly correlative with the Rainbow Gardens Member of Horse Spring Formation) containing clasts (<10 m) of Paleozoic carbonate and Precambrian basement. The conglomerate is intruded by hornblende-quartz monzonite and basalt. These terranes are fault slices that are bounded above and below by low-angle faults and display a reverse stratigraphic order. Precambrian rocks form the structurally highest terrane and Tertiary conglomerate forms the lowest. Geochemical data demonstrate that the intrusive rocks of Tertiary age on Saddle Island are correlative with the Wilson Ridge pluton. Area F. Northern River Mountains Volcanic Section: Basaltic Flows, Dacite Domes, Flows, and Maﬁc Enclaves Park at the Longview Scenic Wayside located approximately 14.5 km (9 miles) east of the Lake Mead National Recreation Area entrance station on Lake Mead Drive. Walk south across the highway to the old Lake Shore Drive now part of the River Mountains Trail System. Walk northwest on old Lake Shore Drive to the ﬁrst dirt road; turn left (south). This 3 km hike provides a view of the diverse volcanic section in the northern River Mountains. Named the Powerline Road volcanic section by Smith (1984), the section includes lower Powerline Road dacite ﬂows and breccia intruded by numerous basalt dikes and rhyolite domes; middle Powerline Road andesite and basalt, and upper Powerline Road rhyolite domes and pyroclastic rocks. See Figures 10 and 15 for stop locations. Stop 13. The hike starts at the base of the section volcaniclastic rocks intruded by plagioclase-biotite dacite of the lower Powerline Road. Numerous dikes of pyroxene, olivine basalt cut this section. Rhyolite dome of the upper Powerline Road just to the west of the road contain ﬂow-banded rhyolite and a marginal

18

Honn and Smith

114o48.000' W

36o05.000' N

114o49.000' W

Area E 10 upper plate

9 36o04.000' N

detachment fault

N

0

MN

.5 0

.5

Figure 14. Topographic map with locations of stops for area (E) Saddle Island. Stops 9 and 10: 9—detachment fault outcrop; 10—outcrop of upper plate rocks.

1 mi 1 km

12.5

O

vitrophyre. The road passes directly through the center of a ﬂowbanded rhyolite dome. The rhyolite contains quartz and sanidine as well as sparse biotite. Stop 14. After passing through the dome, the trail (now a wash) turns to the west. The stratigraphic section to the north of the wash contains middle Powerline Road pyroxene-olivine basalt and agglomerate overlain by rhyolite breccia shed from nearby domes. After 0.4 km the trail again turns to the south. Stop 15. Uphill to the east is a faulted section containing middle Powerline Road basalt (some with clinopyroxene phenocrysts 1–2 cm in size) overlain by pyroclastic surge, ﬂow, and cara-

pace from nearby upper Powerline Road domes. Back in the wash, we will pass the contact of middle Powerline Road basalt with lower Powerline Road dacite, a rhyolite dome (lower Powerline Road), and basalt dikes in lower Powerline Road dacite. Stop 16. Continue walking to the end of the road. Here lower Powerline Road dacite domes are faulted against the purple-colored dacite two (Tdp2) of Smith (1984). This unit contains numerous basalt enclaves some of which have crenulate margins. Originally thought to be a thick dacite ﬂow, recent ﬁeld work indicates that it represents a shallow intrusion that lies stratigraphically above lower Powerline Road volcaniclastic rocks but below lower Powerline Road dacite domes.

Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section

114o51.000' W

36o06.000' N

114o52.000' W

19

Area F 166 rhyolite dome

basalt-rhyolite section

12

Figure 15. Topographic map with locations of stops for area (F) northern River Mountains volcanic section. Stops 11–14: 11—dissected rhyolite dome; 12—basalt-rhyolite contact; 13—pyroclastic surge deposit, dome carapace, and fault; 14—dacite intrusion with maﬁc enclaves.

11 dissected rhyolite dome

basalt-dacite contact

pyroclastic surge deposit, dome carapace, and fault

rhyolite dome and basalt dike

14

36o05.000' N

13

dacite intrusion with mafic enclaves

N

0

MN 12.5

O

.5 0

.5

1 mi 1 km

Stop 17. If time permits, continue walking along the wash to view the interior of a ﬂow-banded rhyolite dome and a basalt dike containing a large block of Precambrian gneiss. Walk back along the wash and return to the Longview Scenic Wayside.

the preliminary 206U/238Pb in zircons by LA-ICP-MS at ETH Zurich. Discussions with Ernie Anderson, Sue Beard, Jim Faulds, Ernie Duebendorfer, Rod Metcalf, Adam Simon, and Terry Spell were very valuable. Sue Beard reviewed the manuscript and provided helpful comments. Lastly, we thank Ernie Duebendorfer for his efforts as co-editor of the ﬁeld guide and his comments on the manuscript.

ACKNOWLEDGMENTS REFERENCES CITED Our understanding of the Wilson Ridge pluton and River Mountains has been inﬂuenced by many people over the last 30 years. We especially thank former University of Nevada, Las Vegas, students Dan Feuerbach, Terry Naumann, Jim Mills, Lance Larson, Deborah Potts, Angela Sewall, Mike Weber, and Ed Eschner, who completed research projects that directly affected our understanding of the area. Adam Simon analyzed

Anderson, R.E., Longwell, C.R., Armstrong, R.L., and Marvin, R.F., 1972, Signiﬁcance of K-Ar ages of Tertiary Rocks from the Lake Mead region, Nevada-Arizona: Geological Society of America Bulletin, v. 83, p. 273– 288, doi: 10.1130/0016-7606(1972)83[273:SOKAOT]2.0.CO;2. Anderson, R.E., Barnhard, T.P., and Snee, L.W., 1994, Roles of plutonism, midcrustal ﬂow, tectonic rafting and horizontal collapse in shaping the Miocene strain ﬁeld of the Lake Mead area, Nevada and Arizona: Tectonics, v. 13, p. 1381–1410.

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Armstrong, R.L., 1966, K-Ar dating using neutron activation for Ar analysis: granitic plutons of the eastern Great Basin, Nevada and Utah: Geochimica et Cosmochimica Acta, v. 30, no. 6, p. 565–600. Armstrong, R.L., 1970, Geochronology of Tertiary igneous rocks, eastern Basin and Range province, western Utah, eastern Nevada, and vicinity, U.S.A: Geochimica et Cosmochimica Acta, v. 34, no. 2, p. 203–232. Duebendorfer, E.M., Sewall, A.J., and Smith, E.I., 1990, The Saddle Island Detachment fault, an evolving shear zone in the Lake Mead area of southern Nevada, in Wernicke, B., ed., Mid-Tertiary extension at the latitude of Las Vegas: Geological Society of America Memoir 176, p. 77–97. Duebendorfer, E.M., Beard, L.S., and Smith, E.I., 1998, Restoration of Tertiary extension in the Lake Mead region, southern Nevada: The role of strikeslip transfer zones, in Faulds, J.E., and Stewart, J.H., eds., Accommodation Zones and Transfer Zones: The regional segmentation of the Basin and Range province: Geological Society of America Special Paper 323, p. 127–148. Faulds, J.E., Smith, E.I., and Gans, P., 1999, Spatial and temporal patterns of magmatism and extension in the Northern Colorado River Extensional Corridor, Nevada and Arizona: A preliminary report, in Faulds, J.E., ed., Cenozoic Geology of the Northern Colorado River Extensional Corridor, Southern Nevada and Northwestern Arizona: Economic implications of regional segmentation structures: Nevada Petroleum Society 1999 Field Trip Guidebook, Reno, Nevada, p. 171–183. Faulds, J.E., Feuerbach, D.L., Miller, C.F., and Smith, E.I., 2001a, Cenozoic evolution of the northern Colorado River extensional corridor, southern Nevada and northwest Arizona: in Erskine, M.C., Faulds, J.E., Bartley, J.M., and Rowley, P.D., eds., The geologic transition, high plateaus to Great Basin: a symposium and ﬁeld guide: the Mackin volume: Paciﬁc Section of the American Association of Petroleum Geologists Publication GB 78 (also Utah Geological Association Publication 30), p. 239–272. Feuerbach, D.L., 1986, Geology of the Wilson Ridge pluton; a mid-Miocene quartz monzonite intrusion in the northern Black Mountains, Mohave County, Arizona, and Clark County, Nevada [M.Sc. thesis]: University of Nevada, Las Vegas, 79 p. Feuerbach, D.L., Smith, E.I., Shaﬁquallah, M., and Damon, P.E., 1991, New K-Ar dates for maﬁc late-Miocene to Pliocene volcanic rocks in the Lake Mead area, Arizona and Nevada: Isochron West, p. 17–20. Feuerbach, D.L., Smith, E.I., Walker, J.D., and Tangeman, J.A., 1993, The role of the mantle during crustal extension: constraints from geochemistry of volcanic rocks in the Lake Mead area, Nevada and Arizona: Geological Society of America Bulletin, v. 105, p. 1561–1575, doi: 10.1130/00167606(1993)105<1561:TROTMD>2.3.CO;2. Honn, D.K., Simon, A.S., Smith, E.I., and Spell, T.L., 2007, The River Mountains volcanic section–Wilson Ridge pluton, a long-lived multiphase midTertiary igneous system in southern Nevada and northwestern Arizona, USA: Eos (Transactions, American geophysical Union), v. 88, p. 52.

Howard, K.A. and John, B.E., 1987, Crustal extension along a rooted system of imbricate low-angle faults: Colorado River extensional corridor, California and Arizona, in Coward, M.P., Dewey, J.F., and Hancock, P.L., eds., Continental Extensional Tectonics: Geological Society of London Special Publication 28, p. 299–311 Larsen, L.L., 1989, The origin of the Wilson Ridge pluton and its enclaves, northwestern Arizona; implications for the generation of a calc-alkaline intermediate pluton in an extensional environment [M.Sc. thesis]: University of Nevada, Las Vegas, 81 p. Larsen, L.L., and Smith, E.I., 1990, Maﬁc enclaves in the Wilson Ridge Pluton, northwestern Arizona: Implications for the generation of a calc-alkaline intermediate pluton in an extensional environment: Journal of Geophysical Research, v. 95, p. 17693–17716. Metcalf, R.V., Smith, E.I., and Mills, J.G., 1993, Magma mixing and commingling in the northern Colorado River extensional corridor; constraints on the production of intermediate magmas—Part I, in Lahren, M.M., Trexler, J.H., Jr., and Spinosa, C., eds., Crustal evolution of the Great Basin and Sierra Nevada: Reno, Nevada, University of Nevada, p. 35–55. Potts, D.A., 2000, Sodium metasomatism and magnesio-riebeckite mineralization in the Wilson Ridge Pluton [M.Sc. thesis]: University of Nevada, Las Vegas, 104 p. Sewall, A.J., 1988, Structure and geochemistry of the upper plate of the Saddle Island Detachment, Lake Mead, Nevada [M.Sc. thesis]: University of Nevada, Las Vegas, 84 p. Smith, E.I., 1982, Geology and geochemistry of the volcanic rocks in the River Mountains, Clark County, Nevada and comparisons with volcanic rocks in nearby areas, in Frost, E.G., and Martin, D.L. eds., Mesozoic-Cenozoic tectonic evolution of the Colorado River Region, California, Arizona and Nevada: San Diego, California, Cordilleran Publishers, p. 41–54. Smith, E.I., 1984, Geologic map of the Boulder City quadrangle, Nevada: Nevada Bureau of Mines and Geology, Map 81, 1:24,000, 1 sheet. Smith, E.I., Feuerbach, D.L., Naumann, T.R., and Mills, J.G., 1990, Mid-Miocene volcanic and plutonic rocks in the Lake Mead area of Nevada and Arizona; Production of intermediate igneous rocks in an extensional environment, in Anderson, J.L., ed., The nature and origin of Cordilleran magmatism: Geological Society of America Memoir 174, p. 169–194. Sun, S.S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the ocean basins: Geological Society of London Special Publication 42, p. 313–345. Weber, M.E., and Smith, E.I., 1987, Structural and geochemical constraints on the reassembly mid-Tertiary volcanoes in the Lake Mead area of southern Nevada: Geology, v. 15, p. 553–556, doi: 10.1130/00917613(1987)15<553:SAGCOT>2.0.CO;2. MANUSCRIPT ACCEPTED BY THE SOCIETY 30 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 11 2008

Late Paleozoic deformation in central and southern Nevada Pat Cashman Jim Trexler Department of Geological Sciences and Engineering, M.S. 172, University of Nevada, Reno, Nevada 89557, USA Walt Snyder Vladimir Davydov Department of Geosciences, Boise State University, Mailstop-1535, 1910 University Drive, Boise, Idaho 83725-1535, USA Wanda Taylor Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154-4010, USA

ABSTRACT In central Nevada, a series of angular unconformities records protracted orogenic activity between middle Mississippian and late Permian time. These unconformities are regional, and can be correlated with lithofacies boundaries at their distal edges. Both the unconformities and the tectonically created sedimentary basins they bound are best expressed in a north-south belt of localities from Winnemucca south to the Las Vegas area. This paper brieﬂy describes seven localities where rocks display both structural and stratigraphic features related to one or more of these unconformities and their related tectonic events. At Edna Mountain, the record is both stratigraphic and structural, and is mostly from the Pennsylvanian. At Carlin Canyon, we will look at both Mississippian and Pennsylvanian folding, thrusting, and unconformities. In the Diamond Range, we will see evidence that Pennsylvanian folding is regionally important. At Secret Canyon, the record is mostly of Permian deformation and sedimentation. In the Hot Creek Range, we will see southern versions of Mississippian stratigraphy, and thrusting that is late Paleozoic in age. In the Timpahute Mountains, complex faulting is also believed to be late Paleozoic. Keywords: Nevada, upper Paleozoic tectonics, unconformities, upper Paleozoic deformation. INTRODUCTION

land basin ﬁll (signaling the collapse of the continental margin) is latest Devonian through mid-Mississippian in age (Poole and Sandberg, 1977). (2) These foreland strata are deformed, and unconformably overlapped by Upper Mississippian through lower Pennsylvanian strata (Silberling et al., 1997). The Pennsylvanian-Permian “Antler overlap sequence,” so named because it unconformably overlies rocks interpreted to have been deformed during the Antler orogeny, also contains angular

We have recognized a complex history of late Paleozoic deformation throughout the Great Basin by focusing on the internal stratigraphy and structure of the “Antler foreland basin” and the “Antler overlap sequence” in central Nevada (Fig. 1). The latest Devonian-Mississippian “Antler foreland basin” comprises a two-part stratigraphy: (1) The initial fore-

Cashman, P., Trexler, J., Snyder, W., Davydov, V., and Taylor, W., 2008, Late Paleozoic deformation in central and southern Nevada, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 21–42, doi: 10.1130/2008.ﬂd011(02). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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unconformities, thus recording several different Pennsylvanian and Permian deformation events. Detailed biostratigraphy enables us to recognize and correlate unconformities between mountain ranges, and is the key to determining the extent and signiﬁcance of late Paleozoic deformation. We have adopted a scheme for naming the late Paleozoic unconformities that is analogous to that used in the Mesozoic of the Four Corners region (Pipiringos and O’Sullivan, 1978; Trexler et al., 2003). Application of this unconformity scheme across much of Nevada reveals that the location of the most intense deformation has changed with time.

In this paper, we summarize the stratigraphic and structural evidence for late Paleozoic deformation at seven Nevada localities (Fig. 2). For each, we present a brief summary of the stratigraphic units involved, a synthesis of the geometry and kinematics recorded by the structures, and the detailed age control that makes unraveling the story possible. The results of previous workers are the basis for many of these summaries. In particular, we acknowledge the work of Silberling et al. (1997) and Tosdal (unpublished mapping) in the northern Pinyon Range, and Dott (1955) in the Adobe Range. Recent thesis research at the University of Nevada, Reno, by Danielle Villa (2007) at Edna Mountain

Edna Mtn. Fm.

Antler Peak

ELY BASIN

4

Battle Cong. Highway Cong. Highway Ls. Iron Point Fm

Figure 1. Tectonostratigraphic units and basins, northeastern and east-central Nevada. Time scale and numerical ages from House and Gradstein (2004), Davydov et al (2004), and Wardlaw et al (2004), all in Gradstein et al. (2004).

Late Paleozoic deformation in central and southern Nevada and Jeremy McHugh (2006) at Luther Waddles Wash in the Hot Creek Mountains was key to understanding these areas. RATIONALE FOR UNCONFORMITY SCHEME Tectonically generated unconformities are stratigraphic boundaries for genetic basin sequences; they are a consequence of the orogenic creation of accommodation space. Orogenic uplift or downwarp is rapid relative to basin response times, and thus these unconformities bracket times when bedrock was eroded and sediment deposited as a response to tectonism. We have proposed the testable hypothesis that tectonically generated unconformities can serve as regional temporal markers (e.g., Snyder et al., 2002; Trexler et al., 2003). We have adopted a scheme of unconformities (Fig. 1); each is deﬁned at a locality where the subjacent strata are deformed by an event not recorded by overlying strata. Two key observations make this scheme work: (1) it is important to choose localities where units above and below the unconformity are as close in age as possible, because successive overprinting of deformation events is expected; and (2) each unconformity is expected to be less angular with distance from the tectonic disturbance, and in even more distal areas the unconformity may be recorded as only a facies change, or not recorded at all. The latter observation implies that one can potentially identify the area of orogenic activity by looking at the intensity of sub-unconformity deformation. SUMMARY OF GEOLOGY AT EDNA MOUNTAIN Stratigraphy of Interest Here Strata at Edna Mountain (Fig. 3) comprise three genetic packages, including (in ascending structural order) the Lower Paleozoic Preble Formation, upper Paleozoic “overlap” carbonate and siliciclastic strata, and the Golconda allochthon. The units of interest for our purposes are in the second group, and include strata deposited on the Preble Formation. These are, from oldest to youngest, the (Morrowan) Iron Point Conglomerate (new name, Villa, 2007) the (Morrowan) Highway Limestone and overlying, undated Highway Conglomerate, the (Virgillian) Antler Peak Formation, and the (Guadalupian) Edna Mountain Formation (Erickson and Marsh 1974b, 1974c). The newly named Iron Point Conglomerate was previously mapped as the Battle Formation (Erickson and Marsh 1974b, 1974c); however, it underlies the Highway Limestone, which is Morrowan in age. The type Battle Formation, near Battle Mountain, contains a thin limestone of Atokan age (Saller and Dickinson, 1982; Theodore, 2000), and clasts of mid-Atokan limestone; it is no older than Atokan. It is therefore 3–5 million years younger than the Iron Point Conglomerate. Unconformities Here All of the stratigraphic contacts at Edna Mountain are unconformities except the gradational contact between the Iron Point

23

Conglomerate and overlying Highway Limestone (Fig. 3). The oldest late Paleozoic unit at Edna Mountain is the Iron Point Conglomerate; it occurs in the upper plate of the Iron Point fault, and its depositional base is not is not preserved here. Iron Point Conglomerate beds grade upward into the Highway Limestone. The upper contact of the Highway Limestone is a karsted, erosional surface (C5), overlain by the “Highway Conglomerate” (Villa et al., 2007), (previously thought to be interbedded with the Highway Limestone [Erikson and Marsh, 1974b, 1974c]). The “Highway Conglomerate” is therefore renamed as a formation in its own right. It occurs in the cores of west-verging synclines in the Highway Limestone, suggesting that deposition was structurally controlled, and that the conglomerate postdates the mid-Pennsylvanian west-vergent folding (Villa et al., 2007; Villa, 2007). There is an angular unconformity between the Highway Limestone and the overlying Antler Peak Formation (C6); the Highway Conglomerate is not preserved at this contact. The Antler Peak also unconformably overlies the folded and metamorphosed Preble Formation along the west ﬂank of Edna Mountain, where the base of the Antler Peak contains clasts of Preble phyllite. The Edna Mountain Formation lies with angular unconformity (P4) on the Antler Peak Limestone and Preble Formation. Structure Several late Paleozoic fold sets, each truncated by one of the unconformities, can be recognized in the “Antler overlap sequence” at Edna Mountain (Fig. 4): West-southwest–verging, asymmetric to overturned folds in the Highway Limestone are unconformably overlain by Antler Peak Formation that does not exhibit this deformation. Open, east-trending folds in the Antler Peak Formation and older units predate the unconformably overlying Edna Mountain Formation. Northeast-trending, southeast-verging folds are developed in the Edna Mountain Formation and all underlying units. Another late Paleozoic structure, the low-angle Iron Point fault, also occurs at Edna Mountain. Originally interpreted to be a thrust fault (Erickson and Marsh 1974b, 1974c), it has been reinterpreted as a top-to-the-northeast low-angle normal fault (Villa, 2007; Villa et al., 2007). Motion along the Iron Point fault post-dates the west-vergent folding and predates the folding around east-trending axes. Implications and Signiﬁcance At least one of the unconformities and deformation events recorded in the “Antler overlap sequence” at Edna Mountain appears to correlate with a similar feature already identiﬁed in Carlin Canyon 100 km to the east (Trexler et al., 2004), thus establishing it as regional in extent. The base of the Iron Point Conglomerate is not preserved at Edna Mountain, so the age and possible correlation of the basal unconformity are constrained only as Morrowan or older (C3 or C2); the C2 unconformity underlies the correlative section at Carlin Canyon. Tight, westverging folding is developed in Morrowan-Atokan rocks at both

24

Cashman et al.

IDAHO

OREGON

nc e

Ra n

ge

Goose Creek Mts.

Tuscarora Mts.

Ind

be

n Ra

Elko Rub y Mt

3

Edna Mountains

2

Carlin Canyon

3

Ferdelford Canyon

4

Diamond Range

5

Secret Canyon

6

Luther Waddles Wash

7

Timpahute Range

nd Ran

ge Ran

ge

inn

Mon

Can yon

itor

Ran

6

Range

reek R

Ran ima Toq u

Hot C

ge

Ran

ange

ge

5

Diamo

Sho

sho

4

Pancak e

ne R

ge

ang

e

Tobi n

Piñon Range

Toi yab e

1

s.

o Ad

Carlin

Field Stop Areas

ge

2

1

Rang

e

Winnemucca

e pe

nde

Snake Mts.

Qu

Tonopah

7

White Mts.

NTS ts. dM oo nw tto Co

lley

Va ens Ow

CA

LI

FO

RN

IA

200 kilometers

200 miles

Figure 2. Regional map with localities discussed in the text.

Las Vegas

Late Paleozoic deformation in central and southern Nevada

Ma

25

Unconformities Here Capt Wo

Guadalupian Cisuralian

Ass

Sak

Art

Kun

Edna Mountain Formation

Vir

Late

Antler Peak Limestone

Mis

Highway Conglomerate

Atok

Des

Mid

C6 C5

Highway Limestone

Early

Pennsylvanian

299

Morrow

P4

Roa

Middle Permian Early

260

Iron Point Conglomerate

318

Iron Point fault

lower Paleozoic

Preble Formation

Figure 3. Stratigraphy at Edna Mountain (see Figure 2 for location). This ﬁgure is from Villa (2007).

Only the boundary between the Tonka and Moleen-Tomera formations is conformable; all the rest, including an unconformity within the Strathearn Formation, are angular unconformities (Fig. 1). The Melandco-Tonka formation boundary is an angular unconformity (C2), particularly well expressed at the west end of the canyon. The Tonka-Moleen-Tomera sequence represents continuous sedimentation from coarse siliciclastics through interbedded carbonate strata. These strata are folded and faulted, and unconformably overlain by the lower (Virgillian) Strathearn Fm (C6) The upper Strathearn (Sakmarian) lies with low-angle discordance (P1) on the lower Strathearn and subjacent strata. The Buckskin Mountain Formation is apparently discordant on the upper Strathearn Formation, here recording the P2 boundary. Structure Three fold sets, at least two of which are truncated by an unconformity and thus unequivocally late Paleozoic, can be recognized in the upper Paleozoic section at Carlin Canyon (Trexler et al., 2004): West-northwest–verging, overturned folds and imbricate thrust faults in the Moleen and Tomera formations are unconformably overlain by the (lower) Strathearn Formation. Open, northeast-plunging folds in the lower Strathearn are erosionally trimmed, then overlain by the upper Strathearn. Northtrending, subhorizontal folds are developed in the upper Strathearn; their age is not constrained here. Implications and Signiﬁcance

Carlin Canyon and Edna Mountain; the oldest undeformed unit overlying these rocks at Edna Mountain is the Antler Peak Limestone. The unconformity below the Antler Peak correlates with the regional C6 unconformity, and constrains the age of west-vergent deformation to at least older than Missourian and younger than Atokan. In addition, we suggest that the unconformity between the Highway Limestone and the Highway Conglomerate is the regional C5 unconformity at the base of the Morrowan; if so, this places a tighter age constraint on the west-vergent deformation. SUMMARY OF GEOLOGY AT CARLIN CANYON Stratigraphy of Interest Here The units of interest here (Fig. 5) are, in ascending order, the Vinini Formation (Roberts Mountains allochthon), the (Osagean-Meramecean) Melandco conglomerate and sandstone, the (Chesterian) Tonka Formation, the (Morrowan-Atokan) Moleen and Tomera formations, the (Missourian-Virgillian) Strathearn Formation(s), and the (Upper Wolfcampian through Leonardian) Buckskin Mountain Formation.

Fortuitous exposure of fossiliferous limestone here provides tight constraints on several of the important tectonic unconformities. The C2 unconformity at the base of the Tonka Formation clearly documents deformation of the initial Antler foreland basin strata, in an angular relationship also seen throughout the Pinyon Range and 100 km to the south in the Diamond Range (see below, and Trexler and Nitchman, 1990; Trexler et al., 1991; Silberling et al., 1997). West-vergent thrusting and overturned folding of Pennsylvanian strata beneath the C6 unconformity records thin-skinned contraction along the continental margin during late Desmoinesian time. Low-angle erosional unconformities P1 and P2 document limited but signiﬁcant deformation in the early Permian in this locality. The structure of this area is documented in detail in Trexler et al. (2004). SIDE EXCURSION TO FERDELFORD CANYON— SUMMARY Stratigraphy of Interest Here Two genetically different packages are preserved at Ferdelford Canyon. The stratigraphy here is the same as at Carlin

469500

470000

470500

471000

471500

471000

471500

469000

4532800 4532000 4531200 4530400 4529600

4529600

4530400

4531200

4532000

4532800

4533600

Cashman et al.

4533600

469000

26

469500

Quaternary units

470000

Iron Point upper plate

Tertiary basalt

Highway Conglomerate

Cretaceous igneous units

Highway Limestone

Golconda allochthon Golconda thrust lower plate

470500

Iron Point Conglomerate Iron Point fault lower plate (Preble Fm)

Edna Mountain Fm

0 Antler Peak Limestone

0.25

0.5

1 Kilometers

Figure 4. Geologic map of Edna Mountain, based on mapping by Erickson and Marsh (1974a, b) and modiﬁed and expanded by Villa (2007).

Permian

Pennsylvanian

Mississippian

Devonian

Wolfcampian

Late

Middle

Kung.

Artin.

Assel. Sak.

Mis.

Vir.

Atokan Des.

Morrowan

Chesterian

Early

Late

Early

Late

Kind.

Osagean Meramecian

4508

Webb Fm.

Melandco Fm.

C2

Tonka Fm.

C3*

Moleen Fm.

C4*

Tomera Fm.

C5*

C6

lower Strathearn Fm.

P1

upper Strathearn Fm.

P2

Buckskin Mountain Fm.

C2

Mm

4509 08

73

Qu

Mw

51

72

Mt

lower to middle Missourian

72

70

Qal

5 83

Ts

48

53

45

Mt

82

60

50

37

49

Qu

27

at Carlin Canyon

40

77

P2

Qal

Qal

Mt

52

65

PPsl

52

8

Psu

7

50 57

60

74

61

44

43

80

PPsl

66

P1

40

44

40

32

33

PPsl

24

48

Qal

24

44

46

40

42 50

85

61

70

45

30

29

24

76

30

44

46

43

53

48

40

C6

46

60

48

23

15

65

42

17

N

Pe

38

54

0

47

51

48

45

45

43

29

34

54

35

80

meters

500

35

9

20

48

14

21

74

78 55

26

50

38

36

35

28

71

74

33

32

82

48

52

77

58

28

29

35

54

85

Ts

1000

68

81

25

85 63

51

42

34

28

86

uppermost Asselian or lower Sakmarian

P2

73

28

38

50

45

45

5 86

45

16

30

54

Psu

52

35

20

20

24

P1

61

Pe

25

late Chesterian or early Morrowan

early Morrowan

77

20

39

45

upper Asselian - lower Sakmarian

late Morrowan

35

42

43

49

51

36

44

34

39

Pbm

Hum boldt R.

63

P1 40

22

Age control: Asselian-Sakmarian Missourian-Virgilian Atokan Morrowan

C3

30

early Atokan

46

40

34

35

5 85

upper Asselian - lower Sakmarian

66

36

15

upper Asselian - lower Sakmarian

32

P2

22

middle Missourian

30

33

Qu

60 53

46

Pe

upper Asselian - lower Sakmarian

54 middle-upper Virgilian

48

18

upper Asselian to lower 43 Sakmarian

5 84

middle-upper Virgilian

middle to upper Virgilian

39

28

middle-upper Virgilian

50

46

PPsl

60

lower to middle Missourian83

C6

55

P1 Ts

PPsl

57

31

middle to upper Virgilian

Key to fold axial surfaces (see Fig. 4 for ages): post-Permian Asselian Desmoinesian-Missourian *missing or conformable

Mw

Mm

Mt

Pe

Pe

PPsl

Psu

Pbm

Ts

Qu

Qal

Figure 5. Geologic map and stratigraphy at Carlin Canyon and Ferdelford Canyon in the Northern Adobe Range (see Figure 2 for location). Map is from Trexler et al. (2004).

Frasnian Famennian

Leonardian

Quaternary undifferentiated Tertiary sedimentary rocks

Quaternary alluvium

Late Paleozoic deformation in central and southern Nevada 27

28

Cashman et al.

Canyon (Fig. 5). Older stratigraphy includes limestone of the Devonian Devil’s Gate Formation, upper Devonian Webb Formation shale and argillite and lower-middle Mississippian Melandco Formation sandstone, shale, and thin conglomerates. The Webb is Famennian in age (Smith and Ketner, 1978) and the Melandco is not dated here. These units are folded and, in some localities, thrust faulted (Jansma and Speed, 1993, Silberling et al., 1997, Tosdal, unpublished mapping). Overlying this folded stratigraphy is the subhorizontal Tonka Formation conglomerate and sandstone, compositionally and texturally similar to the Tonka Formation at Carlin Canyon. The Tonka Formation is dated here as Chesterian, based on macrofossils (Smith and Ketner, 1978). Unconformities Here Two signiﬁcant unconformities are preserved here: the C1 unconformity between the passive-margin Devonian carbonates and the synorogenic, lower Mississippian clastics, and the C2, middle-upper Mississippian unconformity. Strata of the Melandco and Webb formations are deformed and moderately to steeply dipping below the angular C2 unconformity overlain by gently dipping Tonka Formation. Structure Detailed mapping of the Melandco and Webb formations in Ferdelford Canyon documents east-verging folding and thrust faulting (Jansma and Speed, 1993; Silberling et al., 1997; Tosdal, unpublished mapping). Although angular relationships at the C2 unconformity are preserved in several ranges, this locality is the most accessible and best documented example of the deformational style and kinematics of the mid-Mississippian deformation (Silberling et al., 1997; Trexler et al., 2003). Implications and Signiﬁcance Sedimentation in the initial Antler foreland is documented here by the Woodruff and Webb formations (upper Devonian through lowest Mississippian strata) lying unconformably (C1) on strata of the middle Paleozoic miogeocline (e.g., the Devonian Devil’s Gate Formation). The shales and argillites of the Woodruff and Webb are the best evidence for tectonically driven subsidence of the continental margin related to initial emplacement of the Robert Mountains allochthon (Poole and Sandberg, 1977; Murphy et al., 1984). The synorogenic sediments—like those exposed at Ferdelford Canyon—provide better information about the timing of the Antler orogeny than do the ﬁeld relationships for the age of the Roberts Mountains thrust in most areas. This locality also provides kinematic information about middle Mississippian deformation of Antler foreland strata. This deformation has been recognized throughout the Pinyon Range (e.g., Johnson and Pendergast, 1981).

SUMMARY OF GEOLOGY AT THE DIAMOND MOUNTAINS Stratigraphy of Interest Here At Three-Mile Canyon in the Diamond Mountains, the stratigraphy consists of the lower-middle Mississippian Dale Canyon Formation, the upper Mississippian Diamond Peak Formation, the Pennsylvanian Ely Formation, and unnamed Permian units Artinskian through Guadalupian in age (Fig. 6) (Larson and Riva, 1963). The contact (C3) between the Diamond Peak and Ely formations is gradational. Upper Diamond Peak Formation has interbedded limestones that yield Chesterian microfossils (Trexler and Nitchman, 1990; Trexler et al., 1991). The Ely Formation comprises well-bedded, cyclothemic limestones that span the lower and middle Pennsylvanian (Larson and Riva, 1963). Overlying the Ely carbonates are a series of upper Pennsylvanian and Permian unnamed units, separated by regional unconformities (Larson and Riva, 1963; Van Hofwegen, 1995) (see Fig. 6). Unconformities Here The C2 unconformity separates lower and middle Mississippian Dale Canyon Formation from overlying Diamond Peak Formation (Trexler et al., 2003). This unconformity is angular and is identical to the C2 unconformity in Carlin Canyon. The C3 through C6 unconformities here are either expressed as lithostratigraphic boundaries or are trimmed by younger unconformities. The C3 unconformity here is the lithostratigraphic boundary between the Diamond Peak and Ely formations. Dott (1955) recognized upper and lower members of the Ely Formation; we would correlate the boundary between these members with the C4 unconformity. Van Hofwegen documents lower Strathearn Formation (Missourian-Virgillian) above the Ely, and this boundary is the regional C6 unconformity. The P2 overlap strata here are unnamed Leonardian silty and sandy carbonates. Finally, at the top of the range, Guadalupian sandstone and conglomerate caps the section above the P4 unconformity. The large, overturned syncline that forms the Diamond Range has been interpreted as Mesozoic, but predates the Cretaceous Newark Canyon Formation, preserved near Eureka. Structure Our reconnaissance work at Three-Mile Canyon documents folding in the Ely Formation that is erosionally trimmed at the basal contact (C6) of the overlying strata of lower Strathearn age. Pressure solution cleavage is locally developed in the Ely, but not in the overlying Permian units. Subsequent deformation, including a probable fault-bend fold related to a Mesozoic thrust ramp, overprints the Pennsylvanian deformation, and more structural data are needed before this can be conﬁdently removed. However, preliminary results suggest that the pressure solution cleavage dips steeply east after later deformation is removed; this

115°50' Late Paleozoic deformation in central and southern Nevada

39°50'

29

P4

Pgl-gu

Guadalupian limestone

Pa-f

Wolfcampian to Leonardian limestone

IPe

Ely Fm.

Mdp

Diamond Peak Fm.

Mc Mj

(Chainman shale) Dale Canyon Fm.

Dp

Pilot Shale

0

1 miles

2

P2 C3 C2 C1

Joana Limestone

Figure 6. Geologic map and stratigraphy at Three-Mile Canyon, Diamond Range. Map is from Larson and Riva (1963).

N

30

Cashman et al.

would be consistent with west-vergent deformation. This area is a target of future work by the authors. Implications and Signiﬁcance This area was the ﬁrst locality where upper Paleozoic deformation along the western craton margin was documented using the recognition of folded angular unconformities (Trexler and Nitchman, 1990; Trexler et al., 1991). Here, as elsewhere along the deformation belt in central Nevada, the Roberts Mountains allochthon and the Roberts Mountains thrust are very close-by to the west. The implication is that whatever basement structure caused the Roberts Mountains thrust to cut up to the surface at this longitude has also controlled the location and intensity of subsequent deformation. SUMMARY OF THE GEOLOGY AT SECRET CANYON

rian) Diamond Peak Formation is overlain by a 603 m succession of Lower Permian silty micrite to micritic siltstone, ﬁne-grained sandstone, wackestone and packstone strata (the Carbon Ridge Formation of Nolan et al. [1956]), and ~300 m of Lower to Middle (?) Permian conglomerate. This conglomerate was originally considered to be part of the Cretaceous Newark Canyon Formation (Nolan et al., 1956), but the gradational contact with the underlying Carbon Ridge Formation, and rare fusulinids within limestone lenses in the lowermost portion led Steele (1959) and Bissell (1962) to suggest the conglomerate is part of the Carbon Ridge Formation. However, Nolan and Brew (1971) and Strawson (1981), Schwarz (1987) and Baines et al., (1989) have noted the similarity of the conglomerate with the Garden Valley Formation in the Sulfur Spring Range and suggested that this would be a more appropriate name. Here, we adopt the original name “Carbon Ridge Formation” for the entire ﬁne-grained sequence and “Garden Valley Formation” as an acceptable name for the upper conglomerate unit.

Stratigraphic Framework Unconformities The Secret Canyon succession preserves ~900 m of Mississippian–late Early Permian strata (Fig. 7; only lower 640 m shown). The lower 70 m (base not measured) of Mississippian (Cheste-

The lowest Permian unit, which we call Unit 1, is “upper Strathearn” in age (upper Asselian–lower Sakmarian) and rests on

Figure 7. Stratigraphy at Secret Canyon (see Figure 2 for location).

Late Paleozoic deformation in central and southern Nevada the Diamond Peak conglomerates along the P1 unconformity. The equivalent of Unit 1 in the Sulfur Spring Range is solely Asselian in age. The upper contact of Unit 1 is the P2 unconformity which displays 3 m of erosional relief along 100 m of strike. Ammonoids and conodonts recovered from the lower portion of Unit 2 document the abrupt change from the shallow marine succession of Unit 1 to the relatively deeper water depositional environment of these strata. Units 2 through 4 are a shallowing-upwards succession of latest Sakmarian-Artinskian age (late Wolfcampian–early Leonardian) strata. The base of Unit 5 is the P3 unconformity. Unit 5 is poorly exposed, and consists predominantly of silty micrite/micritic siltstone, micritic, and very ﬁne sandstone, with occasional wackestone, packstone and grainstone event beds. The abundance of Phycosiphon grazing traces and in situ biota consisting mainly of small inarticulate brachiopods is reminiscent of Unit 2. Units 5–7 reﬂect another shallowing-upward succession. The P4? unconformity separates the conglomerates of Unit 8 from the underlying units. We speculate that the P5 unconformity may mark the base of the dominantly conglomerate facies of Unit 8B. Structure The stratigraphic section at Secret Canyon is a homoclinal succession on the east limb of a larger anticline. This fold is part of a complex regional structural conﬁguration reﬂecting both Mesozoic and Cenozoic events. The P1 unconformity is slightly angular suggesting, but not proving, some pre-P1 tilting. No major folds have been documented within the section except immediately subjacent to the P3 unconformity. Implications and Signiﬁcance The unconformities at Secret Canyon record a succession of Late Pennsylvanian to Early Permian tectonic events (e.g., Snyder et al., 1991; Trexler et al., 2004). The Ely Limestone and the lower Strathearn Formation that were seen at Carlin Canyon are entirely missing at Secret Canyon. Just 25 km to the south in the Pancake Range, ~1000 m of Ely is preserved above the Diamond Peak Formation and below the Permian section. The P1 unconformity is thus interpreted to reﬂect tectonically controlled differential uplift and subsidence. The P2 boundary between Units 1 and 2 reﬂects an abrupt change in depositional settings and marks the tectonic initiation of the “Dry Mountain Trough” (Stevens, 1977). The origin of the P3 boundary between Units 4 and 5 is unresolved, but a tectonoeustatic mechanism cannot be ruled out. Because they reﬂect deposition intervals of 400,000 yr or less, the 4th and 5th order sequences of Units 2, 3, and 4 are interpreted to be eustatic in origin. The P4 unconformity correlates with the generation of the Phosphoria and Park City basins. The uppermost part of the Garden Valley Formation in the Sulfur Spring Range has yielded Smithian (mid-Early Triassic) conodonts, thus the Garden Valley (Unit 8) at Secret Canyon may host another unconformity that correlates with TR1; this is the subject of an ongoing investigation.

31

SUMMARY OF THE GEOLOGY AT LUTHER WADDLES WASH, HOT CREEK MOUNTAINS Stratigraphy of Interest Here Strata in this part of the Hot Creek Range (Fig. 8) occur in four thrust sheets, each of which contains a distinctive (but as yet unnamed) Mississippian section and therefore records a different part of the Mississippian continental margin (McHugh, 2006). (1) The Big Cow (westernmost) thrust sheet (Fig. 9) contains an Ordovician-Devonian section and fault-bounded early Mississippian and Permian limestones; other parts of the upper Paleozoic section are not exposed. The Mississippian limestones contain graded and scoured beds interpreted as clastic (quartz sand-bearing) and carbonate turbidites. The Permian section includes limestones and dolomites. The Big Cow thrust is unconformably overlain by Early Triassic clastic rocks comprising a basal conglomerate and overlying shale, siltstone and quartz arenite (Fig. 9). (2) The Fishhook Ridge thrust sheet (Fig. 9) contains Devonian carbonates overlain by the Woodruff Formation; the latter represents the Antler foreland basin. This section is overlain by an unnamed, mid-Mississippian (Visean) limestone which is a bioclastic grainstone. It is unusual because other mid-Mississippian rocks in the region (e.g., the Melandco and Dale Canyon formations) are clastic. There is also a fossiliferous coarse-grained Pennsylvanian limestone in this thrust plate; it has been displaced along a later low-angle extensional fault, and its original stratigraphic context is not known. (3) The footwall to both of these thrusts is siliciﬁed and hydrothermally altered Pennsylvanian-Permian crystalline limestone of the White Horse thrust sheet (Fig. 9). (4) The structurally lowest thrust sheet, the Orange Lichen footwall (Fig. 9), is a late Mississippian (Chesterian) turbidite section containing both siliciclastic (both quartz- and chert-clast–bearing) and carbonate turbidites. It matches the Eleana Formation on the Nevada Test Site ~250 km to the south in both composition and age, and is the farthest north this unit is known to occur. Unconformities Here We have not done detailed stratigraphic, sedimentologic or biostratigraphic work on these rocks, so we do not know how many, or which, upper Paleozoic unconformities are preserved here. The age control to date is primarily from conodonts and macrofossils, and much of it is unpublished (see Tables 1–4 in McHugh, 2006, for a compilation and references). Structure Two thrust faults, one east-directed (i.e., foreland-vergent) and the other northwest-directed (i.e., hinterland-vergent) place older Paleozoic rocks over the same Pennsylvanian and Permian footwall section in this area. Superposed folds demonstrate that the east-directed thrusting was the ﬁrst of these two deformation events; stratigraphic relationships bracket it between Permian and

32

Cashman et al.

Big Cow thrust sheet sheet Q- Quaternary alluvial cover

White Horse thrust sheet

Fishhook Ridge thrust sheet

Orange Lichen footwall

STRATIGRAPHIC UNITS:

Q- Quaternary alluvial cover

Q- Quaternary alluvial cover

Q- Quaternary alluvial cover

Siltstone of Big Cow Canyon (TR )

Tv- Tertiary volcanic cover

Tv- Tertiary volcanic cover

Tv- Tertiary volcanic cover

Tv- Tertiary volcanic cover

silstone, sandstone, packstone

light to medium grey limestone and dolostone

Triassic

Mesozoic

Triassic

Mesozoic

Triassic

Mesozoic

Triassic

Mesozoic

Permian carbonate undef. (Pcu)

Penn.-Permian Sed. Rocks (PIPs) silicified crystalline limestone

1

250 Ma

TR

~60 m

5

TR

~60 m

7

TR

~60 m

400 Ma

~100 m

Ddg

457 m

Dnu

544 m

Dd

244 m

Dnl

586 m

SOd

430 m

Oh

91 m

Permian

Permian

1100 m

Dw

~100 m

Ddg

457 m

Dnu

544 m

Dd

244 m

Dnl

586 m

SOd

430 m

Oh

91 m

8

Pl

150 m

9 Mlu

~750 m

Dw

~100 m

10 Ddg

457 m

11 Dnu

544 m

Dd

244 m

13 Dnl

586 m

SOd

430 m

Oh

91 m

12

Pennsylvanian

Pennsylvanian

300 m

MDe

Miss limestone undefined (Mlu)

Mississippian

Dw

Ml

distincively fossiliferous, coarse-grained limestone

fine-grained, crinoidal grainstone

430 m

Paleozoic

300 m Paleozoic

Ml

PIPs

Devonian

Devonian

Paleozoic

350 Ma

3

Devonian

Mississippian

300 Ma

Mississippian

Pennsylvanian

Pennsylvanian

6

Mississippian

300 m

Paleozoic

Pcu

Devonian

2

Permian

Permian

Penn. limestone (Pl)

Miss. limestone (Ml) pale grey bioclastic limestone

Eleana Fm. (MDe) 14 MDe

1100 m

mudrock, litharenite, and chert-pebble conglomerate +/- limestone interbeds

Woodruff Fm. (Dw) Dw

~100 m

Ddg

457 m

Dnu

544 m

Dd

244 m

Dnl

586 m

SOd

430 m

Oh

91 m

mudrock, shale, and chert

Devils Gate limestone (Ddg) siltly, dolomitic lime-mudstone and grainstone

Upper Nevada Group (Dnu)

Oa

Oe

61 m

Oa

305 m

Oe

61 m

Oa

305 m

Silurian

Silurian

Silurian

61 m 305 m

Ordovician

Oe 4

Ordovician

Ordovician

450 Ma

Ordovician

Silurian

dolostone with minor limestone

Denay limestone (Dd)

Oe

61 m

Oa

305 m

fine-grained platy mudstone and packstone with silty horizons

Lower Nevada Group (Dnl) fine-grained, light brown, (where fresh) crystalline dolostone

Silurian-Ord. dolomite (SOd) light grey, recrystallized dolostone

exposed at surface

Hanson Creek dolomite (Oh)

Biostratigraphic age control 1. 2. 3. 4. 5. 6.

Early Triassic Ammoniods (G. Klapper) Permian conodonts (A. Harris) Early Mississippian conodonts (G. Klapper) Ordovician conodonts (R.J. Ross, Jr.) Mesozoic? Ammoniods (E.H. Yockelson) Penn.-Triassic (likely Penn.)conodonts (A. Harris) 7. Early Triassic conodonts (A. Harris)

light grey, recessive weathering dolostone

8. 9. 10. 11.

Pennsylvanian conodonts (M. Kurka) Osage-Meramec conodonts (M. Kurka) mid-Devonian brachiopods (H.W. Dodge, Jr.) Early-mid Devonian crinoids and brachiopods (J.G. Johnson and J.T. Dutro, Jr.) 12. middle Devonian brachiopods (J.G. Johnson) 13. Early Devonian brachiopods (J.T. Dutro, Jr.) 14. Early to Late Mississippian conodonts (E.H. Yockelson)

Eureka quartzite (Oe) light colored, med. to fine-grained, well-cemented qtz arenite

Antelope Valley limestone (Oa) locally fossiliferous wackestone and grainstone

Figure 8. Stratigraphy at Luther Waddle’s Wash, Hot Creek Range (see Figure 2 for location) from McHugh (2006).

Late Paleozoic deformation in central and southern Nevada

33

Figure 9. Geologic map of the area at Luther Waddle’s Wash, Hot Creek Range from McHugh (2006).

Early Triassic. A structurally lower east-directed thrust fault juxtaposes two very different Mississippian units, placing mid-Mississippian limestone over Late Mississippian clastic and carbonate turbidites of the Eleana Formation. An additional, poorly exposed, thrust fault in the northern part of the area appears to place Pennsylvanian-Permian limestone and conglomerate over Early Triassic rocks, documenting a later, post-Early Triassic thrusting event.

Paleozoic deformation, and (2) point out the post-Permian units and structures that cover and disturb the Late Paleozoic unconformities and structures (cf. Fig. 10). The latter point highlights why it has been difﬁcult to identify evidence of Late Paleozoic deformation and unconformities in the region. A thorough understanding of the post-Permian tectonic and depositional history greatly aids in unraveling the Late Paleozoic history because it is important to generate an appropriate paleogeography.

Implications and Signiﬁcance Stratigraphy Our work has documented a signiﬁcant east-directed thrust fault of late Paleozoic age in Luther Waddles Wash; it is overprinted by a northwest-directed thrust fault that may also be late Paleozoic in age (McHugh et al., 2003; McHugh, 2006). Several unusual upper Paleozoic units in this area record depositional environments that were originally west of the Eleana submarine fan system (Trexler et al., 1996; Trexler and Cashman 1997), and are prime targets for further study. SOUTHERN NEVADA On this part of the trip, we will travel from Nye County through Lincoln County and end up in Clark County. The foci of this part of the trip are to: (1) consider the Schoﬁeld Pass fault and whether it is the southern continuation of the belt of Late

The general stratigraphy in northwestern Lincoln County and the central part of eastern Nye County is shown in Figure 11. The Mississippian Joana Limestone (which here includes an unnamed limestone) was deposited before (or during) the Antler orogeny and lies below unconformity C1 (Fig. 1). The unit called Chainman Shale here lacks age control. It is possible that it was deposited in either the Antler foreland basin below C2 or the Antler successor basin above C2. This Chainman Shale is unconformably overlain by the Mississippian Scotty Wash Quartzite. The Scotty Wash Quartzite was deposited between unconformities C2 and C3 in the Antler successor basin. The Ely Limestone, which is regionally considered to be Pennsylvanian in age (Langenheim and Larson, 1973), was deposited in the Ely/Highway basin, most likely between unconformities C3 and C5.

Cashman et al.

P a n c a ke Range

34

117°

River Valley

Qa

ID

CN TB

NV

White

G Ra ran ng t e

y Vall e

UT

?

Qa

road

114°

Winnemuca Belt

Fig. area Sevier Orogenic Belt

CA

N

LC TS

38O

Qa

AZ

G ar

den

Vall e

y

Rail

Tvl

Tv

R.

n yo ane C inn ang Qu R

Tertiary intrusion

Sea

r (S an Vall d Sprin ey g)

man

Tertiary units Cretaceous intrusion

Triassic Upper Paleozoic units, undivided ST

MIT

S

LT

Pen

oye

uPzc

MIR

Pennsylvanian-Permian

PzZc

Lower Paleozoic & Upper Precambrian units, undivided

Tvl

alle at V

Devonian Silurian

y

Cambrian-Ordivician

Tvu

ey

all oV

Qa

bo

PzZc

a Tik

Tvl

Pahranagat R.

nag

Timpahute Range

hra

Pa

Mississippian

115O

Th rus t

Sh

ee

pR

an ge

37O

A Gass

akLas Vegas Pe Range

Las V Shea egas Va lley r Zon e

0

5

10

15 miles

Spring Mountains

115O 30'

Las Vegas

115O

Figure 10. Geologic map highlighting Mesozoic folds and thrust faults of the central Nevada thrust belt (CNTB) and Neogene strike-slip and normal faults. LT—Lincoln thrust, MIR—Mount Irish Range, MIT— Mount Irish thrust, S—Schoﬁeld Pass fault, ST—Schoﬁeld thrust, ID— Idaho, NV—Nevada, UT—Utah, CA—California, AZ—Arizona.

Late Paleozoic deformation in central and southern Nevada

35

Mt. Irish - Golden Gate Thrust Plate Pahranagat Tuff Upper Shingle Pass Tuff Tuff of Hancock Summit Lower Shingle Pass Tuff Monotony Tuff

4000

Ely Limestone Scotty Wash Quartzite Chainman Shale Joana Limestone West Range Limestone

DEVONIAN

Simonson Dolomite Oxyoke Canyon Sandstone

3000 SILUR.

Sevy Dolomite Laketown Dolomite

0

Oxyoke Canyon Sandstone

1000

Laketown Dolomite Fishhaven Dolomite Eureka Quartzite Pogonip Group

0

Schoﬁeld Pass Fault The Schoﬁeld Pass fault, located in the western Timpahute Range, is poorly exposed, but strikes ~N-S. The straight trace of the Schoﬁeld Pass fault across topography indicates a steep dip, but poor exposure prevents determination of the dip direction. Its presence is indicated the stratigraphic differences noticeable near the road to the tungsten mine, which it subparallels. Contact metamorphosed Mississippian rocks on the west side of the road are juxtaposed against Cambrian or Ordovician rocks on the east (Fig. 12). Farther south, Pennsylvanian rocks lie to the west of the fault. The total stratigraphic separation is ~4000 m. The units west of the fault dip ~45°–75°E and are cut and contact metamorphosed by the Lincoln stock. Paleomagnetic data show that this stock is not tilted within the 10° of uncertainty on the data (Taylor et al., 2000). U/Pb dates on zircon from the Lincoln stock yielded a poor age of 98.1 ± 80.8 Ma (Taylor et al., 2000). Therefore the motion on the Schoﬁeld Pass fault that tilted the E-dipping western block occurred after deposition of the Pennsylvanian Ely Limestone and before intrusion of the Cretaceous Lincoln stock. This age permits the speculation that the Schoﬁeld Pass fault is a Late Paleozoic thrust fault. The Schoﬁeld Pass fault also has a history of Cenozoic activity. The Schoﬁeld Pass fault appears to terminate to the

Undifferentiated

SIL

Sevy Dolomite

Goodwin Limestone Windfall Limestone

Simonson Dolomite

ORDOVICIAN

1000

Kanosh Shale Shingle Limestone Parker Spring Fm.

Guilmette Formation 2000

ORDOVICIAN

1500

Fishhaven Dolomite Eureka Quartzite Antelope Valley Limestone

Pogonip Group

2000

CAMB.

2500

Figure 11. Stratigraphy in northwestern Lincoln County. The thrust plates are part of the Jurassic-Cretaceous (?) central Nevada thrust belt, and thus, post date the Late Paleozoic structures. The Rimrock-Freiberg-Lincoln thrust plate lies west of the Mount Irish–Golden Gate thrust plate. Figure from Taylor et al. (2000).

CAMB

3000

Guilmette Formation

Tert. / Quat. alluvium

5000

meters 3500

Quaternary alluvium TERTIARY

v

MISSISSIPPIAN PENN

meters

DEVONIAN

Rimrock - Freiberg - Lincoln Thrust Plate

north at the Penoyer Spring fault. The Penoyer Spring fault is an ~E-W–striking high-angle normal fault that dropped the northern block down relative to the southern block, which contains the Schoﬁeld Pass fault. In contrast, on the southern side of the Timpahute Range, an E-W–striking down-on-the-south normal fault (Crescent Spring fault) terminates at the Schoﬁeld Pass fault. These relationships can be interpreted to indicate that at least the southern part of the Schoﬁeld Pass fault acts as a transfer fault for displacement along the Crescent Spring fault. Several east-striking Cenozoic normal faults that probably post-date early slip on the Schoﬁeld Pass fault transfer slip onto it, reactivating it as a strike-slip fault (Fig. 13). We therefore infer at least some post-mid Cretaceous and post-intrusion slip across the fault. CONCLUSIONS Utilizing the documented late Paleozoic deformation throughout central Nevada, we are now addressing the problem of geographic distribution and regional kinematic relationships. Based on what we know now, the following relationships seem clear: • The mid-Mississippian deformation is expressed most dramatically in the northern Piñon Range; an angular unconformity marks this time period in the Adobe Range

36

Cashman et al. and Diamond Mountains. Farther south in the Nevada Test Site area, and to the east at the longitude of Elko and Ely, Nevada, the unconformity becomes conformable and the signal is stratigraphic. • The mid-Pennsylvanian deformation is well developed at Carlin Canyon, where it includes northwest-vergent folding and imbricate thrusting. At Edna Mountain, well to the west, the mid-Pennsylvanian deformation is characterized by west-southwest–vergent overturned folding. This deformation is also expressed as tight folding at least as far south as the Diamond Mountains. • The earliest Permian deformation is expressed as open folding at Carlin Canyon and becomes more penetrative farther to the south. Our reconnaissance also suggests that these unconformities and the upper Paleozoic deformation history can be traced southeast into geology described by Stevens and Stone (2002) in the Inyo Mountains.

Qa

25

65

Pe

Mj

52 Ms

45 55

79

Mc

Pe

8

63

26 50

Mc Mj

58

Qa 22

10

Mj

59

C Kqm

Op

Cb

28

Mc 78 68

48

Ms

Qa

Kqm Kqm

Qa

78 66

Kqm

75

66 63

60 85

Pe

Qa

65 87 67

37 37 00” 1

115 37 30”

0.5 1

50

45

Qa

22 48

35

64

Qa 0

0.5

This ﬁeld trip begins in Reno, Nevada, but the descriptive road log below begins east of Winnemucca Nevada. For this road log, drive east of Winnemucca on Interstate 80 to the “Golconda” exit where Eden Valley Road and the Midas Road (#789) go northwest and northeast, respectively.

Op

73

33

75 62

Day 1: Reno to Elko

Schofield PassCFault Qa

70

ROAD LOG

TQa

1 MILE

0

1 KILOMETER

Figure 12. Map of the relations across the Schoﬁeld Pass fault. Cb— Bonanza King Formation, Mj—Joana Limestone. Mc—Chainman Shale, Ms—Scotty Wash Quartzite, lPe—Ely Limestone. Kqm— Quartz Monzonite, TQa and Qa—alluvial deposits.

N

Golconda Exit from Interstate 80 (~458 617E/4533 054N) Reset trip odometers here. The prominent outcrops on the hill immediately south of the highway are rocks of the Cambrian (?) Osgood Mountain Quartzite. North and east, in the southern Osgood Mountains and northern Edna Mountain, the quartzite is overlain by the Cambrian Preble Formation.

Reactivated Mt. Irish Fault Crescent Spring Fault System

Lincoln duplex

D D

U

U

Mt. Irish / Golden Gate thrust plate D

Lincoln / Freiberg thrust plate D U

U

ish . Ir ll Mt otwa fo

Tempiute Ridge Block Penoyer Spring Fault System

Reactivated Schofield Pass Fault

Figure 13. Schematic block diagram showing the Cenozoic Penoyer Spring and Crescent Spring fault systems, the offset Mesozoic thrust plates and the reactivated Mount Irish and Schoﬁeld Pass faults. These four fault systems act to limit extensional deformation within the central block by transferring slip onto these nearly orthogonal fault systems.

Late Paleozoic deformation in central and southern Nevada Directly ahead at Edna Mountain the Golconda allochthon rests on the Pennsylvanian-Permian overlap sequence (Highway Limestone, Antler Peak Limestone, Edna Mountain Formation) and the Preble Formation (based on mapping by Erickson and Marsh, 1974b, 1974c). The gray-colored outcrops are the Antler Peak Limestone. The dark outcrops to the southeast along the west side of the range are excellent exposures of greenstone pillow lavas of the Golconda allochthon. Trace element analysis of these and similar rocks suggest they are ocean ﬂoor tholeiites. The following mileage is from the Golconda overpass. Mileage 4.2

4.4 8.4

9.6

17.5

22.6

Description Edna Mountain Formation (4.2 mi from Golconda; ~465 287E/4530 925N). Road cut through the Edna Mountain Formation (Guadalupian); these are immediately subjacent to the Golconda allochthon. Golconda Thrust. Trace of the Golconda thrust crosses the highway—note the excellent exposure (!). Donner Party. Note the Antler Peak Limestone to the north of highway. Just north of here (at Iron Point) is where the Donner Party made their fatal mistake during their trek to California. Here is where the one capable leader was forced to leave the wagon train and his wife and child. The dastardly fellow who began the ﬁght happened to be named “Snyder”—no relation to one of the authors of this guide, of course! Without that hero along, the wagon train spent too many days resting in Reno and was caught by snowstorms while attempting to cross the Sierra Nevada. Happy ending: it was the exiled hero who returned to rescue the stranded wagon train—survivors included his wife and child. Ranch Road exit: Edna Mountain ﬁeld stops. See the guidebook text (above) for descriptions of stratigraphy, unconformities, and structures. Stonehouse exit. Valmy Formation (Ordovician), part of the Roberts Mountains allochthon, makes up the hill immediately adjacent to the highway. Low hills in the distance to the south and the mountain to the southwest (Buffalo Mountain) are underlain by the Havallah sequence of the Golconda allochthon. The large mine north of the Highway is the Lone Tree gold mine. The power plant 3 mi northeast of the highway is the 500 megawatt Valmy Plant operated by Sierra Paciﬁc; it burns coal shipped by train from Utah. Valmy exit (overpass). Battle Mountain (Galena Range, Antler Peak) dominates the skyline to the south. The eastern half of this range consists of the Devonian Harmony Formation (previously considered Cambrian). The Ordovician (and younger?) Valmy Formation and the Devonian Scott Canyon Formation comprise the Roberts Mountains allochthon in the Galena Range. The upper Paleozoic

37

overlap sequence (Battle Conglomerate, Antler Peak Limestone) rests unconformably on the Roberts Mountains allochthon. The Battle Conglomerate here and the Iron Point Conglomerate at Edna Mountain are the oldest overlap to lie depositionally on the allochthon, and thus bracket the age of emplacement as no younger than lower Pennsylvanian. The Golconda allochthon, soled by the Golconda thrust, comprises the western portion of the range and overrides the overlap sequence. ~37.0 Battle Mountain exit (~504 663E/4498 209N). This is the central, Broad Street exit; overpass location. Battle Mountain derives its name from an Indian battle that took place in its northeastern foothills in 1861. The battle began near present-day Beowawe when Shoshone Indians attacked a California-bound wagon train at Gravelly Ford. Following the initial attack, the emigrants regrouped and chased the Indians westward. The Indians took a stand in the rocks near the site of the letters “B.M.” on the Galena Range and were defeated. A settlement named “Reese River Siding” was established ~2 mi north of the battle site at the conﬂuence of the Reese River with the Humboldt River when the Central Paciﬁc Railroad came through this region in 1869; this siding was renamed “Battle Mountain” in January, 1870. The foothills immediate to the south of this exit along Nevada highway 305 are underlain by welded tuff that is the type locality for “ignispumite” (check out your American Geological Institute Glossary)— “a type of rhyolite characterized by lenticles and banding…deposited as an acid, foam lava…transitional with true ignimbrite.” (Ignispumite sounds more like what happens when you have had one too many glasses of Picon punch!) ~88.0 Carlin. This railroad town became a mining town when large gold deposits were discovered near here. These gold mines are on what is now known as the Carlin Trend. Very large open-pit and underground operations are located both north and south of here. Most of these are owned and operated by either Newmont or Barrick corporations. ~94.2 Carlin Canyon: Field trip stops. Refer to guidebook description of the geology here. Continue to Elko, Nevada. Day 2: Elko to Eureka The following distances are from the Intersection of Nevada 278 and Tomera Ranch Road at the southern edge of the town of Carlin, just before (north of) the bridge over the Humboldt River. Stop: Piñon Range, Ferdelford Canyon (Optional) See the discussion of this locality in the guidebook text.

38 Mileage 22.2

37.8 33.5

42.0

Cashman et al. Description Blackburn Oil Field. On the left (east), you can see the remains of the Blackburn oil ﬁeld. The following description of this productive oil ﬁeld was supplied by Jerry Walker of Reno, Nevada: Blackburn Field was discovered in March 1982 by Amoco during the drilling of its #3 Blackburn well (SW/4 SW/4 of Section 8, T27N, R52E, Eureka County). The prospect was based on the nearby Bruffey oil seep, gravity and soilgas anomalies, a poorly deﬁned seismic structural closure, and previous unsuccessful drilling. Initial production from the upper dolomite of the Devonian Nevada Formation at ~7100 feet averaged 79 barrels of oil and 141 barrels of water per day during April 1982. In seven and a half years of work at Blackburn, Amoco drilled four producers and six dry development wells. The ﬁeld was sold to the Petroleum Corporation of Nevada (PETCON) in August 1989. PETCON drilled four wells in Blackburn Field, completing three as oil producers and one as a dry development well. Two of their wells have each produced over a thousand barrels of oil per day. Cumulative production at Blackburn through 2006 was ~5.2 million barrels of oil. Alpha-Tonkin Road intersection. Roberts Mountains overlook (optional stop) (~567 824E/4436 588N). To the southwest, the high mountain range is Roberts Mountains, namesake and type locality for the Roberts Mountains allochthon and the Roberts Mountains thrust fault. The Roberts Mountains allochthon forms the eastern slopes of the range and the Roberts Mountains thrust is exposed just west of the crest of the range, overlying a complete succession of lower Paleozoic miogeoclinal units. The Roberts Mountains thrust is generally deﬁned as the low-angle suture between the autochthonous miogeocline and the overlying Roberts Mountains allochthon. However, while this suture does exist, the timing of emplacement is enigmatic; most localities are clearly reactivated. Sulfur Springs Range view (~570 721E/4425 356N). Look to the east, southeast; the closest prominent ridge is underlain by the Middle-Late (?) Permian Garden Valley Formation. Starting a few miles to the north, a thick band of Permian strata follow (and form) the western edge of the Sulfur Springs Range for ~20 mi. Here, Asselian limestones unconformably overly folded chert of the Vinini Formation (Roberts Mountains Allochthon), followed by a thin Artinskian succession of ammonoid-bearing deeper-water strata. Stratigraphically above these are marginal marine,

45.8

54.8

56.0

57.0 73.04 73.0 ~88.0 ~91.0 0.0

~4.0

~24.0

mixed siliciclastics and carbonates and then ﬂuvial-alluvial sandstones and conglomerates of the Garden Valley Formation. The latter unit is the bluff and cliff former here. McCloud Section (~571 748E/4417 513N). Here, a difﬁcult-to-spot dirt road turns east from the highway, the Bailey Pass Road, leading to the McCloud Spring Section. The entire succession described at mile 42.0 occurs here and was studied in detail by Dora Gallegos (Gallegos et al., 1991). Tyrone Gap Section (~574 425E/4404 228N). A dirt track to the north takes you to the Tyrone Gap Section. Travel ~0.7 mi from the turnoff at Nevada 278, and you will encounter a fork in dirt road, bear right. At ~1.1 mi turn right onto a faint track up the hill toward the section; at 1.2 mi, vehicles can be parked. From here you can walk down section a bit to ﬁnd the contact with the Vinini, but the main traverse will be up the hill through the section described at mile 42.0 (Sulfur Spring Range view). Garden Valley Formation. Here the conglomerates of the Garden Valley Formation intersect the highway. These are probably an alluvial fan sequence of Kungurian or younger age. Sadler Brown Road. Intersection of Nevada 278 and Sadler Brown Road (to north-northeast). Intersection Nevada 278 and U.S. Highway 50. Intersection Highway 50 and Nevada 46 (EurekaThompson Road); to east-northeast. Intersection with U.S. 50. Continue straight on U.S. 50 to Eureka. Eureka. Reset trip odometers in Eureka. North end of Eureka. Reset trip odometer to 0 on U.S. 50 on the north side of town at the bottom of the hill, where a paved road goes right in a narrow “Y.” Follow this paved road out of the canyon and north along the east side of Diamond Valley. Intersection with Diamond Foothill Road. At this complex intersection, follow the main paved road northeast. This is Diamond Foothill Road. Follow the best road northeast (it turns to gravel), with some jogs, to where it gets much closer to the Diamond Range foothills. Where the road runs on the toes of the alluvial fans, watch for a ranch and irrigation circle on the left, at about mile 24. Three-Mile Canyon entrance. Go beyond the canyon mouth, passing a ranch (no residents but active) on the left. 0.3 mi past the ranch is a fenceline road that goes east a short way, then doubles back behind the fenced paddocks to enter ThreeMile Canyon. This road steadily deteriorates, but can be driven ~3 mi into the canyon, to a sheepherders’ trailer and camp. Drive as far as you are willing, and park where you can.

Late Paleozoic deformation in central and southern Nevada Field Trip Stops in the Diamond Range, Three-Mile Canyon Walk to UTM coordinates (approximate) 11S 601535E, 4411422N. See the guidebook text (above) for descriptions of the geology here. Return to Eureka.

~42

~43 ~50

39

Intersection: The valley narrows and the road loops west to a “T” intersection. Ponds and pastures are to the northwest. Turn south (left). Intersection: “T” intersection. Turn left (south) to follow Hicks Station Wash. Luther Waddles Wash Road turns sharply (~150°) right. (568422E, 4291546N).

Day 3: Eureka to Las Vegas Day 3, Part 2: Luther Waddles Wash to Las Vegas Reset trip odometers to 0.0 in Eureka. Mileage 0.0

Description

Bullion Street (~589 427E/4373 580 N). The following distances are measured from the intersection of U.S. Highway 50 and Bullion Street on the south end of the town of Eureka (intersection is near the slag pile; Bullion is to the right (west)). Proceed south on U.S. Highway 50. 9.6 Duckwater Road (~593 593E/4359 666N). Intersection of U.S. Highway 50 and Nevada 379 (Duckwater Road). Turn south onto the Duckwater Road and proceed ~1 mi. 10.8 Windfall Mine Road (~592 713E/4358 040 N). Turn west onto the Windfall Mine Road—this leads directly into Secret Canyon. 12.5 Secret Canyon (~591 148E/4359 892N). The Secret Canyon succession preserves ~657 m of Mississippian–late Early Permian strata. Please see the guidebook text for detailed descriptions. Return to the Duckwater Road, and reset odometer to 0.0 mi. 0.0 Duckwater Road. Drive south along the west ﬂank of Little Smokey Valley. 7.0 Fish Creek Ranch. Here the Duckwater Road turns southeast to cross the valley toward the Pancake Range, and the Duckwater Indian Reservation. Instead, continue southwest. 7.7 “Y” Intersection. Stay to the right on the west side of the valley on the graded road. ~18.2 Landmark: cross a prominent wash here, close to the mountain front on the west. ~25 Landmark: the road crosses three washes here, and a small road leads NW into the range. Continue south. ~26 Cross another, deeper arroyo. ~28 Intersection with a road that goes east and west along this main drainage. Go east on this road ~0.5 mi. ~28.6 Intersection: turn right (south) to continue along the mountain front. Mountains to the west are the Antelope Range. The valley narrows ahead, with the Park Range on your left (east). ~32 Landmark: small road leads west, continue straight. ~37 Intersections: Choose the best road to continue south-southwest. Roads lead both northwest and southeast here.

Return from Luther Waddles Wash along same road. Reset odometer at intersection. Mileage 5.5 11.0

23.0

25.3

Description Turn right (SE) onto Moores Station Road. Turn left (ESE) toward the Pancake Range. Note signiﬁcant incision in basin-ﬁll deposits. The parts of the Pancake Range that can be seen to the NE and SE are composed up Oligocene and Miocene volcanic rocks, mostly ash-ﬂow tuffs. Turn right (SW) onto U.S. 6. Reset odometer. Pause before you turn and look to your right (E). The dark rocks are Pliocene to Quaternary basalt and basaltic andesite of the Lunar Crater volcanic ﬁeld. The rocks to your left (W) are mostly tuffs that were sourced from the Central Nevada Caldera Complex. At 9.5 miles from the last turn, hill on the right (NW) is made up of Quaternary basalt. Morey Peak is the high peak (~10,200 feet, ~3100 m high) on the skyline in the Pancake Range. The peak is composed of the intracaldera tuff of Williams Ridge and Morey Peak (Best et al., 1993). This collapse of this caldera, part of the Central Nevada Caldera Complex, was initiated by the eruption of rhyolitic ash ﬂows that formed the 31.3 Ma Windous Butte Formation (Best et al., 1993). In this region, numerous Oligocene and Miocene out-ﬂow sheets like the Windous Butte Formation that obscure any older structures are visible. At 15.6 miles from the last turn, pass the road to base camp. In the Hot Creek Range, to the NW from here, is the Tybo area. Around 1870, gold was discovered in the Tybo area. The area had its mining boom in the mid-to-late 1870s, and by 1880, it was a ghost town. The rocks in the area are faulted, folded, and locally altered Cambrian to Mississippian marine sedimentary rocks. Turn left (E) onto state highway 375 a.k.a. the Extraterrestrial Highway. Across the valley, the mountains to the SE are the Reveille Range. Most of the visible rocks are Oligocene and Miocene tuffs. This range hosts calderas of the Central Nevada Caldera Complex. The higher topography to the NE is the southern end of the Pancake Range, which is also composed of Oligocene and Miocene volcanic rocks. In the right

40

94.7 100.7 106.7

Cashman et al. light, “basalt” of the southernmost Lunar Crater Volcanic Field may be visible. At 36 miles from the last turn, pass one of the Twin Springs on the left (NE). At 45 miles from the last turn, pass the other of the Twin Springs. Most of the dark-colored rocks near here are Miocene basalt. At 43 miles from the last turn, the major visible valley is Railroad Valley, which is one of two basins in Nevada that has producing oil ﬁelds. At 65 miles from the last turn, the mountains to the NE are the Quinn Canyon Range. This range also is mainly composed of Oligocene and Miocene volcanic rocks. To the north, it contains the Quinn Canyon caldera, which is another part of the Central Nevada Caldera Complex. At 68 miles from the last turn, the stratiﬁed rocks exposed on the lower part of the range to the E (both sides of the highway) are Paleozoic units. The majority of them are Cambrian carbonates. The carbonates are unconformably overlain by Oligocene or Miocene volcanic rocks. At 71.2 miles from the last turn, cross Queen City Summit. Enter Lincoln County. At 74 miles from the last turn, the hill to the NE (~10:00) is made up of Oligocene or Miocene basalt. We are entering Sand Spring Valley (a.k.a. Penoyer Valley). At 84.3 miles from the last turn is the road (right turn) to the Little Al’e Inn in Rachael. This may be an optional stop. The high ridge to the NNE is Tempiute Ridge, which lies on the western end of the Timpahute Range. The ridge contains a slightly faulted east-dipping section of Cambrian to Pennsylvanian units. The high portion of the right is held up by Devonian to Mississippian age marine units. At 91.3 miles, cross over Coyote Summit. Turn left (N) onto Tempiute Mine Road. Schoﬁeld Pass fault ﬁeld stop. Return to NV 375 along same route. Turn left (SE) onto NV 375. Reset odometer. At 2.0 miles from last turn, the hills to the NE (~10:00) are part of the Timpahute Range. They are composed of Oligocene and Miocene tuffs. Normal faults that offset these units are visible from here. The higher mountains behind the hills are composed of Cambrian through Mississippian rocks that are deformed by thrust faults and folds of the Mesozoic Central Nevada thrust belt. At 8.0 miles from last turn, you are in Tikaboo Valley. The range to the NE is the Mount Irish Range and the range to the SE is the Pahranagat Range. The light-colored unit in the near and low part of the Mount Irish Range is the Ordovician Eureka Quartz-

29.5

ite. The thin gray unit below is the Ordovician Antelope Valley Limestone, which is part of the Pogonip Group. Above the Eureka, the section continues into Ordovician Ely Spring Dolomite (dark gray). Above that is the Silurian Laketown Dolomite (a light, dark, light triple stripe). The light gray unit above that is the Devonian Sevy Dolomite. To the east, mediumgray rock of the Devonian Guilmette Formation is faulted down against the Sevy Dolomite. At 11.0 miles from last turn, the gray carbonates and brown sandstones visible in this area are part of the Devonian Guilmette Formation. The quartz-rich sandstones are interpreted to be longitudinal bars. At 16.5 miles from the last turn, the Oligocene and Miocene units near here are mostly sedimentary megabreccias (Jayko, 1990) and the 22.6 Ma tuff of the Pahranagat Formation (Best et al., 1993). At 16.6 miles from the last turn, you approach the Badger Mountain fault in the road cuts. The Badger Mountain fault strikes NW and dips SW. Near the highway, it offsets Oligocene and Miocene units, but farther south it juxtaposes Cambrian and Mississippian rocks. This fault has been interpreted as a thrust fault (Tschanz and Pampeyan, 1970) and a normal fault (Ekren et al., 1977). Jayko (1990) argues that it is a west dipping normal fault with most of its slip occurring prior to deposition of the Oligocene volcanic rocks for four reasons. (1) The hanging wall contains several normal faults that terminate at the Badger Mountain fault. (2) The small syncline visible near Hancock Summit has the correct geometry to be a drag fold related to a down-to-the-SW normal fault. (3) The fault zone contains a large amount of breccia, which is more common along normal faults than thrust faults in nearby areas. (4) The offset of Oligocene to Miocene rocks is minimal compared to the stratigraphic throw of the Paleozoic section. At 22.9 miles from last turn, pass the northern tip of the East Pahranagat Range on the right. In this area is composed mostly of Devonian to Pennsylvanian rocks that are exposed in the limbs of the East Pahranagat Syncline. This fold is one of the structures that was be used to correlate Sevier orogenic belt structures exposed to the south with Central Nevada thrust structures exposed to the north (e.g., Taylor et al., 2000). At 26.0 miles from last turn, to the east across the valley is the Hiko Range. The low pass is the Crystal Wash area. It contains a paleochannel in which the Miocene Hiko Tuff and younger tuffs were deposited (Taylor and Switzer, 2001). This area is also cut by a series of E-W–striking normal faults (Fig. 14). At 29.8 miles, merge with NV 318 from the left. Turn right (SW) onto U.S. 93.

Late Paleozoic deformation in central and southern Nevada

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South E

North E' 6000' Dsi

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Figure 14. Cross section across Crystal Wash and U.S. highway 93 from Taylor and Switzer (2001). Op—Pogonip Group, Oe—Eureka Quartzite, Oes—Ely Springs Dolomite, Sl—Laketown Dolomite, Dse—Sevy Dolomite, Do—Oxyoke Canyon Sandstone, Dsi—Simonson Dolomite, Dg—Guillmette Formation, Th—Hiko Tuff, Tdl— Delamar Lake Tuff, Tsm—Sunﬂower Mountain Tuff, QTa and Qoa—alluvium.

11.95

At 4.8 miles from the last turn, approach community of Ash Springs. Gas station on right (W) and warm springs on left (E). Turn into gas station. Reset odometer. At 4.0 miles from the gas station, note hangingwall anticline in rocks exposed to the west. This anticline lies above the Gass Peak thrust and correlates to a fold we will see along this range farther south (Fig. 10). At 28.5 miles from the gas station, note dissection of the alluvium-lacustrine deposits and old alluvial fan surfaces. The dissection is related to a change in base level for this externally drained valley. This valley (Coyote Valley) drains into Lake Mead and the Colorado River system. The change in base level probably occurred as a result of the opening of the Gulf of California, the associated sea-ﬂoor rifting and the uplift of the Colorado Plateau. At 43.4 miles from the gas station, note folds to the west (right) in Pennsylvanian Bird Spring Formation in the footwall of Mesozoic Gass Peak thrust (Fig. 10). The hanging wall of the thrust is exposed at the skyline and here is composed dominantly of Cambrian and Ordovician carbonate rocks. At 49.3 miles from the gas station, note high angle faults in the rocks to the east (left) in the Arrow Canyon Range. Faults are most visible by looking for displacement in the very dark colored unit, the

97.3

Ordovician Fish Haven Dolomite and the underlying light-colored Ordovician Eureka Quartzite. At 73.5 miles from gas station, take ramp onto Interstate 15 south toward Las Vegas. At 94.7 miles from gas station, exit onto Flamingo Road east. At 95.9 miles from gas station turn left (S) onto Paradise. At 96.5 miles from the gas station, turn left (E) onto Harmon. Arrive at University of Nevada, Las Vegas.

REFERENCES CITED Baines, C.A., Snyder, W.S., and Spinosa, C., 1989, Permian and Cretaceous conglomerates near Eureka, Nevada: Correlation Problems: Geological Society of America Abstracts with Programs, v. 21, no. 5, p. 52. Best, M.G., Scott, R.B., Rowley, P.D., Swadley, W.C., Anderson, R.E., Grommé, C.S., Harding, A.E., Deino, A.L., Christiansen, E.H., Tingey, D.G., and Sullivan, K.R., 1993, Oligocene-Miocene caldera complexes, ash-ﬂow sheets, and tectonism in the central and southeastern Great Basin, in Lahren, M.M., Trexler, J.H., Jr., and Spinosa, C., eds., Crustal Evolution of the Great Basin and Sierra Nevada, Cordilleran and Rocky Mountain Section Field Trip Guidebook, p. 285–311. Bissell, H.J., 1962, Pennsylvanian and Permian rocks of Cordilleran area, in Branson, C.C., ed., Pennsylvanian system in the United States: American Association of Professional Geologists, p. 188–263. Dott, R.H., 1955, Pennsylvanian stratigraphy of the Elko and northern Diamond Ranges, northeast Nevada: American Association of Petroleum Geologists Bulletin, v. 39, p. 2211–2305. Ekren, E.B., Orkild, P.P., Sargent, K.A., and Dixon, G.L., 1977, Geologic map of Tertiary rocks, Lincoln County, Nevada: U.S. Geological Survey Miscellaneous Field Studies Map, I-1041.

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Erickson, R.L., and Marsh, P., 1974a, Paleozoic tectonics in the Edna Mountain Quadrangle, Nevada: U.S. Geological Survey Journal of Research, v. 2, no. 3, p. 331–337. Erickson, R.L., and Marsh, P., 1974b, Geologic map of the Golconda Quadrangle, Humboldt County, Nevada: U.S. Geological Survey Geologic Quadrangle Map GQ-1174, 1:24,000. Erickson, R.L., and Marsh, P., 1974c, Geologic map of the Iron Point Quadrangle, Humboldt County, Nevada: U.S. Geological Survey Geologic Quadrangle Map GQ-1175, 1:24,000. Gallegos, D.M., Snyder, W.S., and Spinosa, C., 1991, Tectonic implications of facies patterns, Lower Permian Dry Mountain trough, east-central Nevada, in Cooper, J., and Stevens, C.H., eds., Paleozoic Paleogeography of the Western United States II: Tulsa, Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, p. 343–356. Gradstein, F.M., Ogg, J.G., and Smith, A.G., 2004, A Geologic Time Scale 2004: Cambridge University Press, 610 p. Jansma, P.E., and Speed, R.C., 1993, Deformation, dewatering, and décollement development in the Antler foreland basin during the Antler orogeny: Geology, v. 21, no. 11, p. 1035–1038, doi: 10.1130/00917613(1993)021<1035:DDADDI>2.3.CO;2. Johnson, J.G., and Pendergast, A., 1981, Timing and mode of emplacement of the Roberts Mountains allochthon, Antler orogeny: Geological Society of America Bulletin, v. 92, p. 648–658, doi: 10.1130/00167606(1981)92<648:TAMOEO>2.0.CO;2. Jayko, A.S., 1990, Shallow crustal deformation in the Pahranagat area, southern Nevada, in Wernicke, B.P., ed., Basin and Range Extensional Tectonics near the Latitude of Las Vegas, Nevada: Geological Society of America Memoir 176, p. 213–236. Langenheim, R.L., Jr., and Larson, E.R., 1973, Correlation of Great Basin stratigraphic units: Nevada Bureau of Mines and Geology Bulletin 72, 36 p., 3 plates. Larson, E.R., and Riva, J.F., 1963, Preliminary geologic map of the Diamond Springs quadrangle, Nevada: Nevada Bureau Mines, scale 1:63,360. McHugh, J.C., 2006, Late Paleozoic contraction in the northern Hot Creek Range, Nye County, Nevada [M.S. thesis]: University of Nevada, Reno, 107 p. McHugh, J.C., Cashman, P.H., Trexler, J.H., Jr., and Taylor, W.J., 2003, Late Paleozoic folding and thrust faulting in the northern Hot Creek Range, Nye County, Nevada: Geological Society of America Abstracts with Programs, v. 36, no. 6, p. 340. Murphy, M.A., Power, J.D., and Johnson, J.G., 1984, Evidence for Late Devonian movement within the Roberts Mountains allochthon, Roberts Mountains, Nevada: Geology, v. 12, no. 1, p. 20–23, doi: 10.1130/00917613(1984)12<20:EFLDMW>2.0.CO;2. Nolan, T.B., and Brew, D.A., 1971, Geologic map of the Eureka quadrangle, Eureka and White Pine Counties, Nevada: U.S. Geological Survey Miscellaneous Geologic Investigations Map I-612. Nolan, T.B., Merriam, C.W., and Williams, J.S., 1956, The stratigraphic section in the vicinity of Eureka, Nevada: U.S. Geological Survey Professional Paper 276, 77 p. Pipiringos, G.N., and O’Sullivan, R.B., 1978, Principal unconformities in Triassic and Jurassic rocks, Western Interior United States—A preliminary survey: U.S. Geological Survey Professional Paper, v. 1035-A, p. A1–A29. Poole, F.G., and Sandberg, C.A., 1977, Mississippian paloegeography and tectonics of the western United States, in Stewart, J.H., Stevens, C.H., and Fritsche, A.E., eds., Paleozoic Paleogeography of the western United States: Tulsa, Oklahoma, SEPM Paciﬁc Section, p. 67–85. Saller, A.H., and Dickinson, W., 1982, Alluvial to marine facies transition in the Antler overlap sequence, Pennsylvanian and Permian of north-central Nevada: Journal of Sedimentary Petrology, v. 52, no. 3, p. 925–940. Schwarz, D.L., 1987, Geology of the Lower Permian Dry Mountain trough, Buck Mountain, Limestone Peak, and Secret Canyon areas, east-central Nevada [M.S. thesis]: Boise State University, 149 p. Silberling, N.J., Nichols, K.M., Trexler, J.H., Jr., Jewell, P.W., and Crosbie, R.A., 1997, Overview of Mississippian depositional and paleotectonic history of the Antler foreland, eastern Nevada and western Utah, in Link, P.K., and Kowalis, B.J., eds., Geological Society of America Fieldtrip Guidebook: Provo, Brigham Young University, p. 161–196. Snyder, W.S., Spinosa, C., and Gallegos, D.M., 1991, Pennsylvanian-Permian tectonism along the western United States continental margin: recognition of a new event, in Raines, G.L., Lisle, R.E., Schafer, R.W., and Wilkinson, W.H., eds., Geology and Ore Deposits of the Great Basin, Symposium Proceedings: Reno, Geological Society of Nevada, p. 5–20.

Snyder, W.S., Trexler, J.H., Jr., and Davydov, V.I., Cashman, P., Schiappa, T.A., and Sweet, D., 2002, Upper Paleozoic tectonostratigraphic framework for the western margin of North America (extended abs.): AAPG Hedberg Research Conference: Late Paleozoic Tectonics and Hydrocarbon Systems of Western North America–The Greater Ancestral Rocky Mountains, p. 58–61. Smith, J.F., and Ketner, K.B., 1978, Geologic map of the Carlin-Piñon Range area, Elko and Eureka Counties, Nevada: U.S. Geological Survey, scale 1:62,500. Steele, G., 1959, Stratigraphic interpretation of the Pennsylvanian-Permian systems of the eastern Great Basin [Ph.D. Dissertation]: Seattle, University of Washington, 294 p. Strawson, F.M., 1981, The geology of the Permian Carbon Ridge Formation, eastcentral Nevada [M.S. thesis]: Reno, Nevada, University of Nevada, 140 p. Stevens, C.A., 1977, Permian depositional provinces and tectonics, western United States, in Stewart, J.H., Stevens, C.H., and Fritsche, A.E., eds., Paleozoic paleogeography of the western United States: Society of Economic Paleontologists and Mineralogists, Paciﬁc Section, Paciﬁc Coast Paleogeography Symposium 1, p. 113–136. Stevens, C.H., and Stone, P., 2002, Correlation of Permian and Triassic deformations in the western Great Basin and eastern Sierra Nevada: Evidence from the northern Inyo Mountains near Tinemaha Reservoir, east-central California: Geological Society of America Bulletin, v. 114, no. 10, p. 1210– 1221, doi: 10.1130/0016-7606(2002)114<1210:COPATD>2.0.CO;2. Taylor, W.J., and Switzer, D.D., 2001, Temporal changes in fault strike (to 90°) and extension directions during multiple episodes of extension: An example from eastern Nevada: Geological Society of America Bulletin, v. 113, p. 743–759, doi: 10.1130/0016-7606(2001)113<0743:TCIFST>2.0.CO;2. Taylor, W.J., Bartley, J.M., Martin, M.W., Geissman, J.W., Walker, J.D., Armstrong, P.A., and Fryxell, J.E., 2000, Relations between hinterland and foreland shortening, Sevier orogeny, central North American Cordillera: Tectonics, v. 19, p. 1124–1143, doi: 10.1029/1999TC001141. Theodore, T.G., 2000, Geology of Pluton-Related Gold mineralization at Battle Mountain, Nevada: Tucson, Arizona, Center for Mineral Resources, Monographs in Mineral Resources Science, 271 p. Trexler, J.H., Jr., and Cashman, P.H., 1997, A southern Antler foredeep submarine fan: the Mississippian Eleana Formation, Nevada Test Site: Journal of Sedimentary Research, v. 67, no. 6, p. 1044–1059. Trexler, J.H., Jr., and Nitchman, S.P., 1990, Sequence stratigraphy and evolution of the Antler foreland basin, east-central Nevada: Geology, v. 18, p. 422– 425, doi: 10.1130/0091-7613(1990)018<0422:SSAEOT>2.3.CO;2. Trexler, J.H., Jr., Snyder, W.S., Cashman, P.H., Gallegos, D.M., and Spinosa, C., 1991, Mississippian through Permian orogenesis in eastern Nevada: post-Antler, pre Sonoma tectonics of the western Cordillera, in Cooper, J.D., and Stevens, C.H., eds., Paleozoic Paleogeography of the Western United States II: Tulsa, Paciﬁc Section SEPM, p. 317–330. Trexler, J.H., Jr., Cole, J.C., and Cashman, P.H., 1996, Middle Devonian–Mississippian stratigraphy on and near the Nevada Test Site: Implications for hydrocarbon potential: The American Association of Petroleum Geologists Bulletin, v. 80, no. 8, p. 1736–1762. Trexler, J.H., Jr., Cashman, P.H., Cole, J.C., Snyder, W.S., Tosdal, R.M., and Davydov, V.I., 2003, Widespread Effects of Mid-Mississippian Deformation in the Great Basin of western North America: Geological Society of America Bulletin, v. 115, no. 10, p. 1278–1288, doi: 10.1130/B25176.1. Trexler, J.H., Jr., Cashman, P.H., Snyder, W.S., and Davydov, V.I., 2004, Late Paleozoic tectonism in Nevada; timing, kinematics, and tectonic signiﬁcance: Geological Society of America Bulletin, v. 116, p. 525–538, doi: 10.1130/B25295.1. Tschanz, C.M., and Pampeyan, E.H., 1970, Geology and mineral deposits of Lincoln County, Nevada: Nevada Bureau of Mines and Geology Bulletin 73, 188 p. Van Hofwegen, D., 1995, Tectonic implications of Pennsylvanian and Permian conodont biostratigraphy at selected locations in the Diamond Range, White Pine and Eureka Counties, Nevada [MS thesis]: Idaho State University. Villa, D.E., 2007, Late Paleozoic Deformation at Edna Mountain, Humboldt County, Nevada [MS thesis]: University of Nevada, Reno, 109 p. Villa, D.E., Cashman, P.H., Trexler, J.H.J., Davydov, V.I., and Taylor, W.J., 2007, Late Paleozoic deformation at Edna Mountain, Humboldt County, Nevada [abs.]: Geological Society of America Abstracts with Programs, v. 39, no. 4, p. 44. MANUSCRIPT ACCEPTED BY THE SOCIETY 29 JANUARY 2008 Printed in the USA

The Geological Society of America Field Guide 11 2008

Active tectonics of the eastern California shear zone Kurt L. Frankel*† School of Earth and Atmospheric Sciences, Georgia Institute of Technology, Atlanta, Georgia 30332, USA Allen F. Glazner† Department of Geological Sciences, University of North Carolina, Chapel Hill, North Carolina 27599, USA Eric Kirby† Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania 16802, USA Francis C. Monastero† Geothermal Program Ofﬁce, Naval Air Weapons Station, China Lake, California 93555, USA Michael D. Strane† Department of Geological Sciences, University of North Carolina, Chapel Hill, North Carolina 27599, USA, and William Lettis & Associates, Inc., 27220 Turnberry Lane, Suite 110, Valencia, California 91355, USA Michael E. Oskin† Department of Geological Sciences, University of North Carolina, Chapel Hill, North Carolina 27599, USA Jeffrey R. Unruh† William Lettis & Associates, Inc., 1777 Bothelo Drive, Suite 262, Walnut Creek, California 94596, USA J. Douglas Walker† Department of Geology, University of Kansas, Lawrence, Kansas 66045, USA Sridhar Anandakrishnan Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania 16802, USA John M. Bartley Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112, USA Drew S. Coleman Department of Geological Sciences, University of North Carolina, Chapel Hill, North Carolina 27599, USA James F. Dolan Department of Earth Sciences, University of Southern California, Los Angeles, California 90089, USA Robert C. Finkel Center for Accelerator Mass Spectrometry, Lawrence Livermore National Laboratory, Livermore, California 94550, USA *[email protected] † Field trip leader Frankel, K.L., Glazner, A.F., Kirby, E., Monastero, F.C., Strane, M.D., Oskin, M.E., Unruh, J.R., Walker, J.D., Anandakrishnan, S., Bartley, J.M., Coleman, D.S., Dolan, J.F., Finkel, R.C., Greene, D., Kylander-Clark, A., Morrero, S., Owen, L.A., and Phillips, F., 2008, Active tectonics of the eastern California shear zone, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 43–81, doi: 10.1130/2008.ﬂd011(03). For permission to copy, contact editing@geosociety. org. ©2008 The Geological Society of America. All rights reserved.

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Frankel et al. Dave Greene Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania 16802 USA Andrew Kylander-Clark Department of Geological Sciences, University of North Carolina, Chapel Hill, North Carolina 27599, USA, and Department of Earth Sciences, University of California, Santa Barbara, California 93106, USA

Shasta Marrero Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, USA Lewis A. Owen Department of Geology, University of Cincinnati, Cincinnati, Ohio 45221, USA Fred Phillips Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, USA

ABSTRACT The eastern California shear zone is an important component of the Paciﬁc–North America plate boundary. This region of active, predominantly strike-slip, deformation east of the San Andreas fault extends from the southern Mojave Desert along the east side of the Sierra Nevada and into western Nevada. The eastern California shear zone is thought to accommodate nearly a quarter of relative plate motion between the Pacific and North America plates. Recent studies in the region, utilizing innovative methods ranging from cosmogenic nuclide geochronology, airborne laser swath mapping, and ground penetrating radar to geologic mapping, geochemistry, and U-Pb, 40Ar/39Ar, and (U-Th)/He geochronology, are helping elucidate slip rate and displacement histories for many of the major structures that comprise the eastern California shear zone. This ﬁeld trip includes twelve stops along the Lenwood, Garlock, Owens Valley, and Fish Lake Valley faults, which are some of the primary focus areas for new research. Trip participants will explore a rich record of the spatial and temporal evolution of the eastern California shear zone from 83 Ma to the late Holocene through observations of offset alluvial deposits, lava ﬂows, key stratigraphic markers, and igneous intrusions, all of which are deformed as a result of recurring seismic activity. Discussion will focus on the constancy (or non-constancy) of strain accumulation and release, the function of the Garlock fault in accommodating deformation in the region, total cumulative displacement and timing of offset on faults, the various techniques used to determine fault displacements and slip rates, and the role of the eastern California shear zone as a nascent segment of the Paciﬁc–North America plate boundary. Keywords: faults, neotectonics, earthquakes, slip rates.

INTRODUCTION The temporal and spatial constancy of fault loading and strain release rates is one of the most fundamental, unresolved issues in modern tectonics. In order to understand how strain is distributed in both time and space across plate boundaries slip rate and displacement data must be compared over a wide range of temporal and spatial scales, from very short-term (tens of years) geodetic data to longer-term (thousands to millions of years) geologic data and from individual fault segments (hundreds to thousand

of meters) to entire fault zones (tens to hundreds of kilometers). Such data are critical to unraveling the complex behavior of lithospheric deformation along plate boundaries. The eastern California shear zone is an ideal natural laboratory in which to study the spatial and temporal evolution of active plate boundary fault systems. As such, the region has been the focus of a number of ﬁeld-based studies in recent years, each with the goal of unraveling the spatial and temporal histories of fault displacements and slip rates. Results of this research highlight the importance of the eastern California shear zone as a

Active tectonics of the eastern California shear zone major piece of the Paciﬁc–North America plate boundary evolution puzzle. The eastern California shear zone is an evolving component of the Paciﬁc–North America plate boundary system (e.g., Faulds et al., 2005; Wesnousky, 2005). This region of predominantly right-lateral strike-slip faults is thought to accommodate ~20%–25% of total relative motion between the Paciﬁc and North America plates (Bennett et al., 2003; Dixon et al., 2000, 2003; Dokka and Travis, 1990; Hearn and Humphreys, 1998; Humphreys and Weldon, 1994; McClusky et al., 2001; Thatcher et al., 1999). The area of active deformation extends northward from the eastern end of the Big Bend of the San Andreas fault near Palm Springs for ~500 km through the Mojave Desert and along the western edge of the Basin and Range east of the Sierra Nevada (Fig. 1). In the Mojave Desert, south of the left-lateral Garlock fault, the eastern California shear zone comprises a 100-km-wide network of NNW–trending right-lateral faults. Geodetic data indicate that elastic strain is accumulating across this zone at a rate of 12 ± 2 mm/yr (Savage et al., 1990; Gan et al., 2000; McClusky et al., 2001; Miller et al., 2001; Peltzer et al., 2001). Seismological and paleoseismological data also indicate that this part of the eastern California shear zone releases strain at a relatively rapid rate; portions of several of these faults ruptured during the 1992 moment magnitude (Mw) 7.3 Landers and 1999 Mw 7.1 Hector Mine earthquakes. Moreover, paleoseismologic data indicate that these two earthquakes are part of an ongoing, ≥1000-yr-long seismic cluster (Rockwell et al., 2000). However, such evidence for rapid strain accumulation and release during the recent past is at odds with longer-term slip-rate data, which suggest that the long-term, cumulative slip rate across the Mojave part of the eastern California shear zone is much slower. Recent work in the Mojave section of the eastern California shear zone indicates the total long-term slip rate across this fault system is on the order of 5–7 mm/yr, or about half of the current rate of strain accumulation determined from space-based geodesy (Oskin and Iriondo, 2004; Oskin et al., 2006, 2007) These observations suggest a pronounced strain transient across the Mojave section of the eastern California shear zone. The Garlock fault bisects the eastern California shear zone, forming a major geologic and physiographic boundary between the Mojave Desert and western Basin and Range (Fig. 1). This fault system began accommodating Paciﬁc–North America plate boundary deformation in the middle to late Miocene and has long been recognized as a major tectonic feature in the region (Burbank and Whistler, 1987; Davis and Burchﬁel, 1973; Loomis and Burbank, 1988; Monastero et al., 1997; Smith et al., 2002). Since the late Pleistocene, there has been at least 18 km of offset along the fault, yet little to no signiﬁcant earthquake activity has occurred over at least the past ~300 yr (Carter, 1980; Dawson et al., 2003; McGill and Sieh, 1991). The Garlock fault is somewhat enigmatic in that nowhere does it appear to offset, or be offset by, NW-trending eastern California shear zone faults. A large question in eastern California shear zone tectonics remains as to

45

how Paciﬁc–North America plate boundary strain is transferred through, or around, the Garlock fault. Nonetheless, displacement from the southern part of the eastern California shear zone in the Mojave Desert is funneled northward across the Garlock fault onto four main fault systems; the Owens Valley, Panamint Valley–Hunter Mountain–Saline Valley, Death Valley–Fish Lake Valley, and Stateline fault zones (Fig. 1). The only fault in the northern eastern California shear zone to experience signiﬁcant seismic activity is the Owens Valley fault, which last ruptured in a Mw 7.6(?) earthquake near the town of Lone Pine in 1872 (Beanland and Clark, 1994). A series of down-to-the-NW normal faults transfer slip between the Owens Valley, Panamint Valley–Hunter Mountain–Saline Valley, and Death Valley–Fish Lake Valley faults (Fig. 1; Dixon et al., 1995; Lee et al., 2001a; Reheis and Dixon, 1996). In contrast to the strain transient south of the Garlock, the region-wide rate of dextral shear across these four faults appears to have remained constant at 9–10 mm/yr over late Pleistocene to recent time scales, although large uncertainties in geologic slip rates on the Owens Valley and Panamint Valley–Hunter Mountain–Saline Valley faults make the strain budget somewhat tentative (Bacon and Pezzopane, 2007; Bennett et al., 2003; Frankel et al., 2007a; Lee et al., 2001b; Oswald and Wesnousky, 2002). Farther north, dextral motion between the Sierra Nevada block and North America is focused on two faults bounding the east and west sides of the White Mountains: the White Mountains fault zone to the west and the Fish Lake Valley fault zone to the east. Of these two, global positioning system (GPS) data suggest that the Fish Lake Valley fault system is storing at least 75% of the elastic strain accumulating in this region (Dixon et al., 2000). However, recent late Pleistocene slip rate studies on these two faults reveal a geologic versus geodetic rate discrepancy similar to that in the Mojave, whereby the White Mountains and Fish Lake Valley faults account for less than half of the region-wide rate of shear determined from GPS data (Bennett et al., 2003; Frankel et al., 2007b; Kirby et al., 2006). Until recently, the total displacement on any of the major faults in the eastern California shear zone was thought to be only a few tens of kilometers (Burchﬁel et al., 1987; Guest et al., 2007; Moore and Hopson, 1961; Niemi et al., 2001; Ross, 1962). However, the correlation of Jurassic and Cretaceous dikes, Cretaceous leucogranites, and a Devonian submarine channel across Owens Valley suggests that right-lateral deformation, in what is now the eastern California shear zone, is much greater than this and may have started as early as 83 Ma (Bartley et al., 2008; Glazner et al., 2005; Kylander-Clark, 2003; Kylander-Clark et al., 2005). The total dextral component of shear across Owens Valley serves as an important constraint on the timing, duration, and rate of strikeslip deformation throughout the region. This ﬁeld guide reviews some of the newest results bearing on displacement and slip rate histories for many of the faults in the eastern California shear zone. We begin with a discussion of temporal variations in rates of deformation across the Mojave Desert, then move on to Miocene to recent slip rate variations

46

Frankel et al.

Figure 1. Shaded relief index map of Quaternary faults, roads, towns, and ﬁeld trip stops in the eastern California shear zone. Most faults are from the U.S. Geological Survey Quaternary fault and fold database (http://earthquake.usgs.gov/regional/qfaults). Arrows indicate relative fault motion for strike slip faults. Bar and circle indicates the hanging wall of normal faults. Field trip stop location numbers are tied to site descriptions in the ﬁeld guide section. AHF—Ash Hill fault; ALF—Airport Lake fault; B—Bishop; BF—Blackwater fault; BLF—Bicycle Lake fault; BM—Black Mountains; BP—Big Pine; Br—Baker; Bw—Barstow; By—Beatty; CA—California; CF—Cady fault; CLF—Coyote Lake fault; CoF—Calico fault; CRF—Camp Rock fault; DSF—Deep Springs fault; DV-FLVF—Death Valley–Fish Lake Valley fault; EPF—Emigrant Peak fault; EV— Eureka Valley; FIF—Fort Irwin fault; FM—Funeral Mountains; GF—Garlock fault; GFL—Goldstone Lake fault; GM—Grapevine Mountains; HF—Helendale fault; HLF—Harper Lake fault; HMSVF—Hunter Mountain–Saline Valley fault; I—Independence; LF—Lenwood fault; LLF— Lavic Lake fault; LoF—Lockhart fault; LP—Lone Pine; LuF—Ludlow fault; LV—Las Vegas; M—Mojave; MF—Manix fault; NV—Nevada; O—Olancha; OL—Owens Lake; OVF—Owens Valley fault; P—Pahrump; PF—Pisgah fault; PV—Panamint Valley; PVF—Panamint Valley fault; R—Ridgecrest; S—Shoshone; SAF—San Andreas fault; SDVF—southern Death Valley fault; SLF—Stateline fault; SPLM—Silver Peak–Lone Mountain extensional complex; SNF—Sierra Nevada frontal fault; SP—Silver Peak Range; T—Tonopah; TF—Tiefort Mountain fault; TMF—Tin Mountain fault; TPF—Towne Pass fault; WMF—White Mountains fault; YM—Yucca Mountain.

Active tectonics of the eastern California shear zone along the Garlock fault, followed by a history of faulting in Owens Valley spanning the past 83 m.y., and ﬁnish with a discussion of late Pleistocene slip rates and kinematics of the Death Valley–Fish Lake Valley fault system. The second part of the guidebook serves as a road log to 12 ﬁeld sites that we feel provide an excellent representation of the Late Cretaceous to late Holocene deformational history of the eastern California shear zone. ACTIVITY OF THE EASTERN CALIFORNIA SHEAR ZONE ACROSS THE CENTRAL MOJAVE DESERT: EXAMPLE OF THE LENWOOD FAULT Paleoseismic records from southern California indicate clustering of earthquake activity over millennial time-scales on individual major faults (Garlock: Dawson et al., 2003; San Andreas: Weldon et al., 2004; San Jacinto: Rockwell et al., 2006) and across fault systems (Mojave eastern California shear zone: Rockwell et al., 2000; Los Angeles basin: Dolan et al., 2007). Understanding the origin of such clustered earthquake activity is clearly important for seismic hazard forecasts and also bears on understanding of fault strength and loading processes. Discrepancy between geologic slip rates and short-term loading measured from geodesy can indicate variation in the distribution of loading rate across a system of faults (e.g., Friedrich et al., 2003; Bennett et al., 2004). The Mojave Desert portion of the eastern California shear zone contains well-documented examples of such conditions. The sinistral Garlock fault is undergoing loading at less than half its well-established geologic slip rate of 5–7 mm/yr (McClusky et al., 2001; McGill and Sieh, 1993). Conversely, the Blackwater fault appears to be loading at a rate that is up to an order of magnitude faster than its geologic slip rate (Peltzer et al., 2001; Oskin and Iriondo, 2004). New evidence from a comprehensive slip-rate investigation across the Mojave Desert section of the eastern California shear zone reveals that the geologic versus geodetic discrepancy exists system-wide (Oskin et al., 2006). New slip rates were determined from mapping of high-resolution airborne laser swath mapping (ALSM) topography and dating of offset features along six dextral faults that comprise the eastern California shear zone at ~34.7°N. From east to west these faults are the Helendale, Lenwood, Camp Rock, Calico, Pisgah-Bullion, and Ludlow (Fig. 1). Based on offset alluvial fans and mid-Quaternary basalt ﬂows, the sum of fault slip-rates across the entire province is at most ~6 ± 2 mm/yr (Oskin et al., 2006). This geologic rate is only half the geodetic rate of 12 ± 2 mm/yr across the province (Sauber et al., 1994; Bennett et al., 2003) but is comparable to that expected from the paleoseismology of Rockwell et al. (2000). This rate discrepancy supports the conclusions of Peltzer et al. (2001) and Dolan et al. (2007) that strain accumulation alternates between different components of the southern California fault system, though the length- and time-scale of such alternation remains to be determined. Consistency of strain release rates measured over ~60 k.y. and ~650 k.y. on the Calico fault (Oskin et al., 2007) and the overall consistency between geologic rates and paleo-

47

seismicity suggests that the time-scale of loading variation is shorter than the ~5000 yr return period of Mojave eastern California shear zone earthquake clusters. The Lenwood Fault The Lenwood fault is a typical example of the active dextral faults that transect the Mojave Desert region. Total dextral slip on the northern Lenwood fault is only 1.0 ± 0.2 km, based on offset of the Lenwood anticline (Strane, 2007). This offset is consistent with ~1 km displacement of a 7.33 ± 0.12 Ma (40Ar/39Ar; A. Iriondo, unpublished data) basalt ﬂow in the Fry Mountains by the southern Lenwood fault (Carleton, 1988). South of the Fry Mountains, the Old Woman Springs fault branches westward from the Lenwood fault. Both fault strands terminate southward at the foot of the San Bernardino Mountains. Similar bifurcations occur approaching southern terminations of other Mojave eastern California shear zone dextral faults (Fig. 1). Slip on the northernmost Lenwood fault diminishes into parasitic folding of the north limb of the Lenwood anticline (Dibblee, 1970) before reemerging as the Lockhart fault (Fig. 1). Geologic and geomorphic mapping of a 48 km2 ALSM topographic survey by Strane (2007) deﬁned both Miocene and late Quaternary stratigraphy and displacements. This ﬁeld guide examines two sections of the northern Lenwood fault where Quaternary alluvial fan markers have been deposited across the fault and cut by subsequent dextral slip (Figs. 2 and 3). Alluvial Fan Ages and Stratigraphy Quaternary units are subdivided based upon three weathering-related, age-diagnostic criteria (Wells et al., 1987): (1) development of desert varnish and rubiﬁcation (reddening) on surface clasts, (2) development of desert pavement, and (3) smoothing of depositional bar and swale microtopography. Smoothing over time of depositional alluvial landforms such as bars and channels is especially apparent in comparisons of differently aged surfaces with high resolution topography (e.g., Frankel and Dolan, 2007). In the Mojave Desert region, recognizable bar and swale morphology is generally only present in active washes, but can be preserved in very young alluvial surfaces and is useful for identifying Holocene alluvial fans (Bull, 1991). Dissection also serves to differentiate alluvial surfaces. Generally, the more highly dissected fans are older (Bull, 1991). However, the smoothing effects of soil and pavement development can counteract formation of channels. Based on these criteria, three generations of older alluvial fan deposits are differentiated along the Lenwood fault. The oldest deposit, Q2a, is characterized by a continuous, smooth, well-interlocked desert pavement with angular, darkly varnished and rubiﬁed surface clasts. The intermediate deposit, Q2b, has less well-developed pavement and clast coatings. Q2b can also be differentiated from Q2a by the presence of well-deﬁned, but shallow incised channels and by the degree of soil development. The youngest deposit, Q3, preserves relict bar and swale topog-

48

Frankel et al. 8 OM

980

10 10 8 M O

Park here 990

O M

8

Pull-apart basin

Q2a

Q2a Qal

Le

nw oo

Q2a dF au

lt

1 Q2a

Barstow Formation

60 m wide

2

270 m P

45 ± 5 m Q2b

Q2b

Figure 2. Hillshaded airborne laser swath mapping (ALSM) topography, illuminated from the west (top) and geologic interpretation (bottom) near Stop 1 (Fig. 1), showing stream deﬂections along the Lenwood fault. Gray patches are gaps in ALSM data. One-meter contours (light gray lines) with 10-m index contours (dark lines, elevations labeled). Middle to late Miocene Barstow Formation consists of poorly consolidated conglomerate at this locality. Arrows: 1—deﬂection of stream by ~270 m around Q2b fan. Though modiﬁed by faulting, this deﬂection is interpreted as primarily a result of aggradation of fans derived from the southeast; 2—~60-mwide stream apparently undeﬂected by faulting. Slip since emplacement of Q2b probably does not exceed the width of this stream; 3—Pull-apart basin in Q2b hosts a small channel that is deﬂected 45 ± 5 m. Star symbols show locations of 10 Be sample pit (P) and modern stream samples (upstream of S).

3 Q2b

Qal

North

S

Q2b

500 meters Q3

raphy. Clast coatings are very light and pavement development is mostly absent except for some winnowing of ﬁne materials from relict channels. Cosmogenic dating using in situ accumulation of 10Be in quartz (e.g., Gosse and Phillips, 2001) yielded a very welldeﬁned age for the Q2b surface (Fig. 4). Samples of both small quartz pebbles and sand-sized bulk sediments were extracted from a 1.8-m-deep hand-excavated pit located in a large, wellpreserved Q2b alluvial fan. The sample site is situated within the southern ﬁeld stop (Stop 1; Figs. 1 and 2), where topographic relationships indicated a thick accumulation of Q2b deposits. This site was found to be characterized by very high 10Be inheritance, equivalent to >60 k.y. of surface-exposure. Interestingly, it was also found that this inheritance varied with grain size (Fig. 4). Despite the high inheritance, the depth-proﬁle approach yielded a very well constrained date of 37 ± 7 ka for the abandonment of the Q2b alluvial fan depositional surface.

Fault Offsets and Along-Strike Slip Rate Gradients Offsets of Q2b alluvial fans and related inset channels are apparent in both the southern and northern ﬁeld stops (Stops 1 and 2; Figs. 1, 2, and 3), though the northern set of offsets is better deﬁned. Here, combined constraints from a pair of inset channels and a shutter ridge yield an offset of 30 ± 5 m (Fig. 3). Channel deﬂections in the southern ﬁeld stop (Stop 1) are complicated by multiple fault strands and fan emplacement along the fault scarp (Fig. 2). Offsets of the larger channels proved difﬁcult to constrain, but must be <60 m. One small channel that crosses a pull-apart basin appears offset 45 ± 5 m. However, it is possible that the channel ﬂowed parallel to the Lenwood fault as the pull-apart basin formed, reducing the amount of slip required. If this amount of slip accrued here since abandonment of the Q2b alluvial fan, then the slip rate along this portion of the Lenwood fault would be 1.2 ± 0.2 mm/yr, which

Active tectonics of the eastern California shear zone

49

0

88

920

940 90

960

0

0

90

920

940

Park here 960

Q2a

Q3

Q3

North

Qal Pickhandle Formation

Q2b Barstow Formation

1 29 ± 8 m

2 30 ± 13 m

Pickhandle Formation

Figure 3. Hillshaded airborne laser swath mapping (ALSM) topography, illuminated from the west (top) and geologic interpretation (bottom) near Stop 2 (Fig. 1), showing stream deﬂections along the Lenwood fault. Gray patches are gaps in ALSM data. Ten-meter index contours shown as dark lines with elevations labeled. Early Miocene Pickhandle Formation here consists of consolidated conglomerate with poorly exposed bedding except for occasional, 2–25-m-thick lenses of monolithologic megabreccia. Mid- to late Miocene Barstow Formation consists of well-bedded sandstone and conglomerate at this locality. Arrows: 1 and 2—locations of a pair of channels deﬂected by dextral slip along the Lenwood fault; 3—shutter ridge formed by dextral offset of channel wall and adjacent hillslopes. Combining all of these constraints yields a best-ﬁt offset of 30 ± 5 m since incision of Q2b.

3 30 ± 5 m

500 meters

is 50% higher than the rate of 0.8 ± 0.2 mm/yr calculated from the better constrained northern site. Such along-strike gradients in slip rate have become increasingly recognized in the eastern California shear zone (e.g., Frankel et al., 2007a, 2007b; Oskin et al., 2007) and represent an important source of potential error that needs to be addressed when compiling geologic rate budgets. Summary The Lenwood fault is a typical example of the rather slowmoving (~1 mm/yr) dextral faults that comprise the eastern California shear zone across the central Mojave Desert. Displaced Q2b alluvial fans and related inset channels document well-constrained offset of 30 ± 5 m and a less well-constrained offset of 45 ± 5 m at two sites along the northern Lenwood fault. A cosmogenic depth-proﬁle age of 37 ± 7 ka for a Q2b alluvial fan results in slip rates of 0.8 ± 0.2 mm/yr and 1.2 ± 0.2 mm/yr for these respective sites.

MIDDLE TO LATE MIOCENE STRATIGRAPHY OF THE SUMMIT RANGE AND RED ROCK CANYON: IMPLICATIONS FOR TEMPORAL VARIATIONS IN SLIP RATE ON THE GARLOCK FAULT The Summit Range, or Summit Diggings as it is labeled on some maps, is located west of the Trona–Red Mountain Road, ~20 km south of Ridgecrest on the south side of the Garlock Fault (Stop 3, Fig. 1). It covers a relatively small ~20 km2 area with the western half consisting of late middle Miocene to earliest Pliocene volcanic, volcaniclastic, and sedimentary rocks that rest nonconformably on Cretaceous (86 Ma) quartz monzonite basement (ages given above and in the rest of the sections for the Summit Range and Red Rock Canyon areas are from Monastero and Walker, unpublished data). Dibblee (1967) included the Summit Range in his discussion of the Lava Mountains although the latter area extends signiﬁcantly farther south and east and is generally underlain by younger rocks than in the Summit Range (Smith et al., 2002). Recent work indicates that the Summit Range

50

Frankel et al. 10

Be concentration (atoms/g x 105) 0 2 4 6 8 10 12

0 Q2b Depth (cm)

50

100

150

37 ± 7 ka 34.7433°N 116.9222°W 1014m

pebbles sand

200 modern stream 10

Figure 4. Plot of Be concentration versus depth for quartz-bearing sediment from Q2b fan surface adjacent to the Lenwood fault (P in Fig. 2). Both 1–3-cm-sized pebbles and sand-sized material were measured. Beryllium10 concentrations from these different grain sizes showed a consistent offset from one another, with higher inheritance for sand-sized grains. The order switches for the modern stream sample site (S on Fig. 2), where sand-sized grains have signiﬁcantly lower concentration than shielded samples from a 150 cm depth. Gray bands show 95% conﬁdence of mean concentration of 10Be with depth for independent age models for each grain size. Both models yielded an identical mean age of 37 ka. Normalizing all pit samples to a common inheritance value, and including the sample of pebbles from the modern stream sample, yields an age of 37 ± 7 ka (95% conﬁdence) for Q2b.

rocks constitute a separate volcanic center that can be directly correlated with middle to late Miocene Dove Spring Formation rocks that crop out in the Red Rock Canyon area on the north side of the Garlock fault, ~30 km west-southwest of the Summit Range (Stop 4, Fig. 1). Middle to Late Miocene Stratigraphy of the Summit Range The stratigraphy in the Summit Range is shown in Figure 5. Locally overlying the crystalline basement is gray andesite porphyry. This ﬁne-grained rock has phenocrysts of plagioclase that have altered rims and hornblende that has been completely replaced by limonite. The age for this rock is 15.6 ± 0.5 Ma (from (U-Th)/He on zircon), which makes it equivalent to the early middle Miocene Cudahy Camp unit 5 (Tc5) of the Red Rock Canyon area. Above this unit is a series of deep red to tan arkosic sandstones, conglomerates, olive-drab–brown mudstones, limestones, and cherts. Most of these rocks appear to be water-lain sedimentary rocks with varying amounts of ash, lithic clasts, and pumice. Some of the rocks appear to be lahars with crystals of feldspar and biotite scattered throughout a dense red matrix. Many of the mineral grains are hydrothermally altered (e.g., reaction rims on

feldspars and limonite replacement of maﬁcs). These rocks are probably equivalent to the lower Dove Spring Formation of Red Rock Canyon (Loomis and Burbank, 1988). Overlying and locally interbedded with the sedimentary units is a series of tuffs, lahars, and debris ﬂows that display a wide range of compositions, colors, thicknesses, and degrees of induration. The base of this sequence consists of gray to dark reddish brown to light purple-brown, ash-rich debris ﬂows and lahars (Fig. 5). Some contain volcanic bombs that range in size from 20 cm to 30 cm and are fractured in-place, indicating they were hot when emplaced. The next overlying unit consists of a series of tuffs that vary widely in color and degree of induration (Fig. 5). Most are white, cream, green, or gray-green, probably representing varying amounts of alteration. Some are clearly air-fall tuffs that conform to irregularities, while others appear to be small ash-ﬂow events. There are also instances where apparent epiclastic units occur within the tuff sequence as denoted by the occurrence of rounded grains of quartz and other accessory minerals. This entire unit attains a thickness of >15 m in the Summit Digging location, but the total thickness and inclusion of all tuffs in the sequence at correlative outcrops varies widely. A distinctive pink to orange-pink ash-ﬂow tuff lies within this unit and crops out prominently in the middle of the Summit Range area. This pumiceous lapilli tuff is rich in lithic clasts with compositions including propylitically altered porphyritic extrusive volcanic rocks, plutonic rocks of varying composition, quartz, and banded rhyolite (Fig. 5). It also has a large percentage by volume of pumice clasts. It reaches a thickness of several meters in outcrop in the Summit Digging area, and crops out widely on the ﬂoor of the valley, where it can be found overlying the tan and deep red lahars, debris ﬂows, and sandstones, and in other places be found beneath them. The age of this unit is 11.8 ± 0.9 Ma (from (U-Th)/He on zircon). The next highest stratigraphic unit consists of a series of dacite domes, ﬂows, and tuffs clustered in a 5 km2 area situated on the western and southern margins of the Summit Range. These units overlie the white tuffs that form the valley ﬂoor between the westernmost ﬂows and the easternmost domes and overlie the tuffs on a small, elliptical, NE-trending hill 2 km south of the main diggings. Units in the domes are quite distinctive because of their coarse phenocrysts of twinned (Carlsbad habit) orthoclase that reach lengths of several centimeters. These are embedded in a ﬁne-grained, reddish-brown matrix with biotite and small quartz grains. On the easternmost dome, there is a vertical fabric to the rock. The dome is surrounded on two sides by outcrops of lava ﬂows with the same, only ﬁner-grained, mineral composition as the dome. These ﬂows exhibit abundant ﬂow features, such as alignment of phenocrysts in ﬂow bands, and contorted, brecciated ﬂow fronts. Flows at the southeast corner of the dome dip 30–40° outward and show well-developed cooling column structure. The radiometric age of this dome is 11.0 ± 0.2 Ma (40Ar/39Ar). A second dome, located ~1 km southwest of the dome just described, is petrologically similar although it is highly altered,

Active tectonics of the eastern California shear zone

Summit Range

Dacite domes and associated flows and tuffs 11.7 ± 0.6 Ma to 11.02 ± 0.19 (M&W)

51

Red Rock Canyon

Tda4 air fall tuff 8.5 ± 0.13 Ma (L&B)

Td5

Tda3 air fall tuff 8.4 ± 1.8 Ma (L&B) Tda2 air fall tuff 10.4 ± 1.6 Ma (L&B) Tdb3 basalt flow 10.5 ± 0.25 Ma (L&B)

Td4 Tdb2 basalt flow 11.3 ± 0.3 Ma (M&W)

Td3 White and green tuffs Tda1 lapilli tuff 11.8 ± 0.9 Ma (L&B) (Tuff of Dutch Cleanser Mine)

Lapilli tuff 11.8 ± 0.9 Ma (M&W)

White and green tuffs Air fall tuff 11.7 ± 0.3 Ma (M&W) Red to tan arkosic conglomerate and sandstone,olive drab to brown mudstone, limestone, and sandstone

Td2 Conglomerate, sandstone, mudrock, limestone, chert

Gray andesite porphyry 15.6 ± 0.5 Ma (M&W)

Atolia quartz monzonite 86.0 ± 1.0 Ma (M&W)

Figure 5. Schematic stratigraphic sections for the Summit Range and Red Rock Canyon areas. Important and/or dated units shown in pattern; unpatterned rocks consist of conglomerate, sandstone, mudstone, limestone, and chert. Ages are from Monastero and Walker (unpublished data), denoted M&W, and from Loomis and Burbank (1988) and Whistler and Burbank (1992), denoted L&B. Units and descriptions for the Red Rock Canyon area are from Loomis and Burbank (1988).

mostly massive, coarsely porphyritic dacite. There is pervasive argillic alteration of orthoclase and biotite phenocrysts and a general vertical fabric, which suggests ﬂow within a vent. The heavily altered area covers ~75–100 m2 and stands 8–10 m high. Numerous examples of monolithologic breccias reminiscent of vent structures can be found on the top of the dome. Flows extend to the south and west of the dome complex and thin rapidly to 1–3 m within 1 km. The age of this dome is determined to be 11.7 ± 0.6 Ma (40Ar/39Ar). Middle to Late Miocene Stratigraphy of the Red Rock Canyon Area Approximately 30 km west-southwest of the Summit Range is the Red Rock Canyon area (Stop 4, Fig. 1). Situated at the

southwest end of the El Paso Mountains, this area is the type section for the middle to late Miocene Dove Springs Formation (Loomis, 1984). The stratigraphic succession for this interval is shown in Figure 5. Loomis (1984) divided the rocks into six units, ranging in age from 13.5 Ma to ca. 7 Ma; the Dove Springs rests disconformably on the early to middle Miocene Cudahy Camp Formation. The hiatus between the Cudahy Camp and Dove Springs Formations was determined by Loomis and Burbank (1988) to be 15.1–13.5 m.y. based on magnetostratigraphy, K-Ar, and ﬁssion track dates. The lowest part of the Dove Springs section contains conglomerate and sandstone consisting of poorly sorted, massive to crudely stratiﬁed units dominated by volcanic and plutonic clasts in an ash-rich matrix. The unit grades continuously upward into massive, dirty arkosic sandstones. The top of this unit is

52

Frankel et al.

arbitrarily deﬁned as the point where the sandstones and conglomerates become less massive and more stratiﬁed. There are no age-speciﬁc geochronology markers in this unit. Member 2 of the Dove Spring Formation (Td2) crops out extensively in the south-central part of Red Rock Canyon State Park where it is dominated by conglomerates, sandstones, and volcanic tuffs (Fig. 5). Loomis (1984) determined that the provenance of these rocks was from a source located to the south and east. The sedimentary and epiclastic rocks consist of grains of quartz, feldspars, and biotite, with varying degrees of rounding, and ash, pumice, and lithic clasts from volcanic and plutonic sources. They vary in color from deep red to tan, are intimately interleaved, and constitute between 200 and 400 m of section. In the middle of Td2, there is a pink lapilli ash-ﬂow tuff (Tda) that is a prominent ridge-former in the Red Rock Canyon area (Fig. 5). The unit is made up of two separate ﬂows, with a cooling break between them. The rock is rich in pumice clasts (0.5– 2.0 cm) and lithics (propylitically altered volcanics, banded rhyolites, and plutonic fragments) in an ash- and crystal-rich matrix. This unit grades northeastward to a white to light gray air-fall tuff dated at 11.8 ± 0.9 Ma (Whistler and Burbank, 1992). There are several thin (tens of centimeters) white air-fall and pumicerich tuffs below this unit that have been dated at 11.7 ± 0.6 Ma (40Ar/39Ar). Overlying the pink tuff is a continuation of the Td2type rocks described in the foregoing paragraph. Higher in the section there are two basalt ﬂows (Tdb2 and Tdb3 of Loomis, 1984), the lower of which have been dated at 11.3 ± 0.3 Ma (40Ar/39Ar; Monastero and Walker, unpublished data). These ﬂows are interleaved with more of the Td2 type rocks although they are grouped into higher units of the Dove Spring.

time of initiation of extension in Death Valley. This would mark the beginning of movement on the Garlock fault inasmuch as there would have been no need for such a structure prior to this time. The assumption is, therefore, that initial movement on the eastern part of the Garlock fault began ca. 15 Ma, which means that there was ~30 km of sinistral offset on the Garlock from inception of movement until 12–11 Ma. This translates to a rate of offset of between 7 and 10 mm/yr. Since that time, an additional ~35 km of offset has resulted in the present spatial relationship of the two sites, which calculates to a slip rate of 2.75–3.0 mm/yr—much slower than the rate of the earlier period. The speciﬁc cause of this dramatic decrease in offset rate on the Garlock fault is not known with certainty at present. Lonsdale (1991) and Atwater (1989) document a change in the conﬁguration of the Paciﬁc and North America plates ca. 12.5 Ma. Lonsdale (1991) contends that the plate offshore Baja California stopped spreading at this time, subduction ceased, and the Rivera triple junction jumped southward to the tip of the Baja peninsula. Atwater (1989) found that at this same time there was simultaneous formation of a continuous boundary between these two plates from central California to the tip of Baja. There are numerous examples of volcanic outbursts during this time in the nearby Owlshead Mountains (Calzia and Ramo, 2000), Death Valley (Thompson et al., 1993; Troxel, 1994), and Eagle Crags volcanic ﬁeld (Monastero et al., 1994). It is possible, and actually highly probable, that the rate of sinistral offset on the Garlock fault varied considerably from 12 to 11 Ma until present, but such correlations are left to a more detailed understanding of the stratigraphy of the middle to upper Dove Spring Formation and the volcanic units found in the Lava Mountains (Smith et al., 2002).

Implications of Unit Correlation at Red Rock Canyon and the Summit Range

Summary

Although the relative positions of the Summit Range volcanic center and the southern Red Rock Canyon stratigraphic sequences at the time of emplacement cannot be determined with absolute certainty, it is clear that they correlate in age and lithology over a wide range of time. It is suggested that the Summit Range location was the source area for most of the volcanic components of the lower Dove Spring Formation in Red Rock Canyon. Based on thickness of the pink lapilli tuff and the fact that the Tda unit makes a transition to a pure white air-fall tuff in a north-northeasterly direction, we interpret that the two sites were more or less juxtaposed across what is now the Garlock fault when the volcanic center was active between 12 and 11 Ma. The notion of the Garlock being an intracontinental transform (Davis and Burchﬁel, 1973) that acts as an accommodation zone for extension north of the fault from a relatively unextended area south of the fault, implies that movement on the Garlock could not have initiated before the onset of extension in the southwest Basin and Range. Wernicke et al. (1988) report the onset of extension in the Las Vegas Valley shear zone–Lake Mead area ca. 16 Ma, and McKenna and Hodges (1990) cite 15 Ma as the

The rocks exposed in the Summit Range and Red Rock Canyon demonstrate that at 12–11 Ma the two areas, which are now separated by ~35 km of sinistral offset on the Garlock fault, were directly opposite one another across that structure. This implies that initial motion on the Garlock fault was on the order of 10 mm/yr from ca. 15–12 Ma. If this is correct, then the later offset rate is a factor of three to four times slower, averaged over the past 11 Ma. BEDROCK EVIDENCE FOR 65 km OF DEXTRAL OFFSET ACROSS OWENS VALLEY, CALIFORNIA, SINCE 83 Ma Historic earthquakes (Beanland and Clark, 1994), paleoseismology (Bacon and Pezzopane, 2007; Lee et al., 2001b), and geodesy (Dixon et al., 2003) indicate that the modern tectonic regime of Owens Valley is dominated by right slip that accommodates a signiﬁcant fraction of the relative motion between the North America and Paciﬁc plates. However, the total magnitude and the timing of lateral slip across Owens Valley have been uncertain. Conventional wisdom has long been that net lateral

Active tectonics of the eastern California shear zone LLsz

mapped Cretaceous(?) dextral-transpressive shear zone Offset(?) Markers

GBD

?

0.706

TP

38°

0.706

Sierra Nevada

Devonian channel (Stevens et al., 1997)

MMP

(Kistler, 1993)

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e on

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CA

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lt z fau

Golden Bear and Coso Dikes Moore (1963, 1981) ﬁrst recognized the Golden Bear dike and mapped it for ~15 km, from its western termination in the northern Mount Whitney quadrangle to where it disappears under Owens Valley near Independence (Fig. 6). The dike actually comprises from one to three branching granitic dikes that range from 5 to 30 m thick, and locally have thick (50 cm) cataclastic margins. The Coso dike swarm includes two groups of steeply dipping E-striking dikes that are separated by ~6 km across strike (Dufﬁeld et al., 1980; Whitmarsh, 1998). Major dikes of the Coso swarm are from 5 to 25 m thick, commonly accompanied by thinner (<2 m) subparallel dikes. The Golden Bear and Coso dikes (Stops 6 and 8; Figs. 1 and 6) are porphyritic quartz monzonite with <5 vol% maﬁc minerals, 2–4 cm K-feldspar phenocrysts (25–30 vol%), and equant, euhedral, bipyramidal quartz crystals (~15 vol%) up to 1 cm across. Zircon U-Pb analyses by Kylander-Clark et al. (2005) of two Golden Bear dike samples deﬁned an emplacement age of 83.4 ± 0.4 Ma. Whitmarsh (1998) reported a U-Pb zircon date of ca. 88 Ma (later interpreted as 84 ± 1 Ma; J.D. Walker, 2003, personal commun.) from one Coso dike, and Kylander-Clark et al. (2005) obtained several concordant, or nearly concordant, zircon fractions that cluster between 82 and 86 Ma from two other Coso dike samples. In addition, major-element, trace-element, and isotopic geochemistry of the dikes are compatible with correlation of the Golden Bear and Coso dikes (Kylander-Clark et al., 2005). Although the Coso and Golden Bear dikes range up to 30 m thick, neither continues on strike across Owens Valley (Fig. 6). The Inyo Range east of the Golden Bear dike (Ross, 1965) exposes latest Precambrian through early Paleozoic sedimentary rocks that are intruded by Jurassic and Cretaceous plutons and sparse dikes of the Independence swarm, but none of the dikes resembles the Golden Bear dike in either petrology or dimensions. Similarly, on strike westward from the Coso dikes in the Olancha Peak area of the Sierra Nevada, various Mesozoic

Jurassic plutons (Ross, 1962)

SRF MMP TR

0.706 Sri = 0.706 line

119°

Offset Geologic Markers

Cretaceous Golden Bear/Coso dike sets and wall rocks (Kylander-Clark et al., 2005) Jurassic Independence dike swarm IDS (Bartley et al., 2008)

y alle sV en Ow

slip across the valley is no more than 10 km (Moore and Hopson, 1961; Ross, 1962) and that right slip in Owens Valley began in Plio-Pleistocene time (Lee et al., 2001b). However, recently identiﬁed offsets of several bedrock geologic markers indicate dextral displacement of 65 ± 5 km (Fig. 6), including, from youngest to oldest: (1) the Golden Bear dike in the Sierra Nevada and the Coso dike swarm in the northern Coso Range; (2) 102–103 Ma leucogranite and maﬁc granodiorite plutons into which the Golden Bear and Coso dikes were intruded; (3) the axis of maximum dilation of the 148 Ma Independence dike swarm; and (4) a Devonian submarine channel complex and other geologic features identiﬁed in the Mount Morrison roof pendant in the Sierra Nevada and east of Tinemaha Reservoir at the base of the Inyo Range (Fig. 6). The offsets are interpreted to record the same displacement that must have accumulated since 83 Ma, when the Golden Bear–Coso dikes were emplaced.

53

GBD

N TR SRF SRFsz

IR IDS CD

main figure

LLsz

36°

LL* 118°

CR AR

? IDS

Figure 6. Locations of markers that support 65+ km of dextral offset across Owens Valley. Although correlation of the Tinemaha (TP) and Santa Rita Flat (SRF) plutons suggested little lateral offset (Ross, 1962), plutons of similar age and lithology are widespread on both sides of Owens Valley. Offset markers include (1) the densest part of the Independence dike swarm (Bartley et al., 2008); (2) the Sr 0.706 isopleth (Kistler, 1993); (3) a Devonian submarine channel at Tinemaha Reservoir and in the Mount Morrison pendant (Stevens et al., 1997); (4) the 83 Ma Golden Bear (GBD; Stop 8, Fig. 1) and Coso dike sets (CD; Stop 6, Fig. 1; Kylander-Clark et al., 2005); and (5) the 102–103 Ma plutons that the Golden Bear and Coso dikes intrude (Kylander-Clark et al., 2005). AR—Argus Range; CR—Coso Range; IDS—Independence dike swarm; IR—Inyo Range; LL—Little Lake; LLsz—Little Lake shear zone of Bartley et al. (2008; Stop 5, Fig. 1); MMP—Mount Morrison pendant; SRFsz—Santa Rita Flat shear zone of Vines (1999); TR—Tinemaha Reservoir; WM—White Mountains; WMsz—White Mountain shear zone of Sullivan and Law (2007).

plutons are present, but no thick granite porphyry dikes (Diggles, 1987; Diggles et al., 1987). Early Cretaceous Plutons The Golden Bear dike intrudes Cretaceous plutons and minor Mesozoic metavolcanic and metasedimentary rocks (Moore, 1963, 1981; Chen and Moore, 1982; Kylander-Clark et al., 2005; Saleeby et al., 1990). Except near its western terminus, the dike

54

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mainly intrudes coarse leucogranite of the 102 Ma Bullfrog and Independence plutons and granodiorite of the 103 Ma Dragon pluton (Saleeby et al., 1990; Kylander-Clark et al., 2005). Granitic plutons dated ca. 102 Ma are common farther west in the Sierra Nevada (e.g., intrusive suite of Yosemite Valley; Ratajeski et al., 2001), but the Bullfrog, Independence, and Dragon are the only plutons of this age known at the eastern range front of the Sierra Nevada. The Coso dikes intrude Sierran-like Mesozoic plutons and minor Mesozoic metasedimentary and metavolcanic rocks. Plutonic wall rocks include the leucogranite of Cactus Flat, which yielded a U-Pb zircon date of 101.6 ± 0.8 Ma (Kylander-Clark et al., 2005), making it the only pluton east of Owens Valley to yield an Early Cretaceous age. The leucogranite of Cactus Flat is commingled with undated granodiorite that resembles the Dragon pluton.

(1998) to propose that rocks in these two areas are remnants of a major Devonian submarine channel that has been dextrally offset ~65 km on a cryptic fault, termed the Tinemaha fault, in northern Owens Valley (Fig. 7). This interpretation is also compatible with ~65 km of dextral offset of the initial 87Sr/86Sr = 0.706 isotopic isopleth proposed by Kistler (1993) in the Mono Lake area. The Tinemaha fault was originally considered to be Middle Triassic in age, based on the supposition that the fault was intruded by the Late Triassic Wheeler Crest Granodiorite (Stevens et al., 1998). However, given the present evidence for 65 km of well-dated Late Cretaceous or younger dextral displacement in the southern Owens Valley, it now seems clear that the fault in the northern Owens Valley must also be Cretaceous or younger in age (Stevens et al., 2003).

Independence Dike Swarm Dikes of the Independence swarm vary widely in abundance, and some parts of the swarm are compositionally diverse, whereas others are uniformly maﬁc (e.g., Carl and Glazner, 2002; Fig. 7). The internal anatomy of the swarm thus was mapped in search of previously unrecognized tectonic offsets. Dike thicknesses and orientations were recorded along 84 cross-strike traverses in areas of good to excellent exposure on both sides of Owens Valley (Bartley et al., 2008; Fig. 7). Dilation by intrusion of the Independence dike swarm ranges from 0 to 44%, but dilation values >5% are restricted to distinct, relatively narrow zones in the eastern Sierra Nevada near Independence, the northern Alabama Hills, and the southern Coso and Argus ranges and Spangler Hills (Fig. 7). In high-dilation areas, the swarm is also lithologically diverse, including maﬁc, intermediate, and felsic dikes. Elsewhere in the swarm where dilation is less (mainly <1%), only maﬁc dikes generally are present. The most southerly high dilation values on the western side of Owens Valley were found in the northern Alabama Hills. The nearest high dilation area on the eastern side of the valley is ~75 km to the southeast in the southwestern Coso Range. This offset is uncertain by at least ±10 km because the dike swarm is diffuse and intersects the valley at a low angle (~30°).

Net Offset Each of the above correlations requires right-lateral displacement of a similar magnitude. The displacement required by the Golden Bear–Coso dike correlation is insensitive to the precise location of the fault(s) that accommodated it because the dikes dip steeply and strike at high angles to the valley. Restoring the easternmost Golden Bear outcrops to a position adjacent to the westernmost Coso dike outcrops yields ~65 km of right slip (Fig. 6). Aligning the Golden Bear dike with the northernmost mapped exposures of the Coso dikes reduces the estimate to ~60 km. Correlations of the 102–103 Ma plutons, the maximum dilation axis of the Independence dike swarm, and the Devonian channel deposits are also compatible with 60–65 km of right-lateral offset.

Devonian Submarine Channel and Other Geologic Features On McGee Mountain in the Mount Morrison pendant, a distinctive facies of the Middle Devonian Mount Morrison Sandstone consists of quartz sandstone and coarse conglomerate containing clasts of chert and limestone with fragments of tabulate coral (Greene and Stevens, 2002; Stevens and Greene, 1999). The unit ﬁlls a channel cut deeply into older rocks and is interpreted as the feeder channel of a major submarine fan complex. A lithologically identical conglomerate, also of Middle Devonian age and containing fragments of tabulate corals, is exposed east of Tinemaha Reservoir at the base of the northern Inyo Mountains, 65 km to the southeast of McGee Mountain. The lithologic and structural similarities between these areas led Stevens et al.

Discussion

Timing of Offset Correlation of the Golden Bear and Coso dikes requires that the 65 km offset accumulated since ca. 83 Ma. At current estimates of long-term slip rate across Owens Valley of 2–3 mm/yr, 65 km of offset could accumulate in ~20–30 m.y. This is compatible with the late Cenozoic right-oblique subduction setting of California (Atwater and Stock, 1998), and slip could have commenced when the San Andreas system started to form in the late Oligocene. Indeed, there is evidence that dextral slip began in the Mojave Desert in the early Miocene (Glazner et al., 2002). However, most stratigraphic, structural, and thermochronologic data suggest that the modern dextral slip regime commenced in the Pliocene (Lee et al., 2001b; Monastero et al., 2002; Stockli et al., 2003). To accumulate 65 km of offset since then would require a slip rate of >20 mm/yr, which is more than half of the total Paciﬁc–North America plate motion and almost certainly is incompatible with regional constraints. Among the problems with this hypothesis is that the Garlock fault shows little evidence for disruption by N- to NW-striking faults. Many dextral faults in the Mojave Desert do, indeed, die out before reaching the Garlock fault, and it is not clear how ongoing dextral slip south of the

Active tectonics of the eastern California shear zone

119°W

A

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ins ounta e M Whit ens Ow

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ada

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Argus

#dikes/km (B) 0-10

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%dilation (A) 0-2%

SH

50 km

Figure 7. Map of quantitative properties of the Jurassic Independence dike swarm, modiﬁed from Bartley et al. (2008). Shaded areas are preCenozoic bedrock. Symbols are centered on midpoints of traverses. Dashed symbols indicate traverses where observed dikes may be Cretaceous and therefore not part of the Independence swarm (see Bartley et al., 2008, for further discussion). (A) Percent dilation by diking. Note offset of high-dilation zone between the Alabama Hills (AH) and Coso Range and the distinct northern boundary on both sides of Owens Valley where high dilations drop to zero. SH—Spangler Hills. (B) Number of dikes crossed per kilometer. The pattern resembles that in (A), but traverses in the southeastern Sierra Nevada (dashed) appear more prominent owing to much higher abundances of thin (5–30 cm) maﬁc dikes. Such a high proportion of thin maﬁc dikes is atypical of the Independence swarm and more characteristic of Cretaceous dikes in the region. (C) Kamb contour plot of poles to dikes measured on traverses. GF—Garlock fault; SAF—San Andreas fault; CA—California; NV—Nevada.

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Frankel et al.

Garlock fault is transmitted to the north (e.g., Oskin and Iriondo, 2004). Thus, it seems most likely that a signiﬁcant fraction of the total 65 km of slip—perhaps as much as 50–60 km—is signiﬁcantly older, possibly Late Cretaceous (Bartley et al., 2008). Laramide-age right slip along the eastern margin of the Sierra Nevada may have been linked to a south-directed extensional detachment system in the southern Sierra Nevada (Wood and Saleeby, 1997) that unroofed high-pressure rocks in latest Cretaceous–early Tertiary time. Dextral shear of this age would predate formation of the Garlock fault (Monastero et al., 1997) and thus eliminate that conﬂict. Locus of Slip The fault zone responsible for offset of the various markers lies between the Sierra Nevada and White-Inyo-Coso Ranges (Fig. 6). Although mostly buried by alluvium, it can be located within 100 m at Little Lake on the west side of the Coso Range, where Jurassic plutonic rocks with abundant Independence dikes are juxtaposed against a distinctive Jurassic orthogneiss that lacks such dikes (Stop 5; Figs. 1 and 8; Bartley et al., 2008). Cretaceous and Jurassic intrusive rocks on both sides of the fault in this area are cut by transpressive ductile-brittle shear zones with a persistent dextral component of shear. These shear zones, and similar Late Cretaceous dextral transpressive shear zones exposed in the Inyo (Vines, 1999) and White (Sullivan and Law, 2007) Mountains, probably belong to a family of structures that accommodated 50+ km of dextral offset in Late Cretaceous– early Tertiary time. Approaching the main fault contact at Little Lake, cataclastic microstructures increasingly overprint the ductile microstructures in the shear zones. This spatial pattern may reﬂect the reactivation of a Late Cretaceous shear system by the modern right-slip fault zone. Summary Several independent geologic markers indicate 65 ± 5 km of net dextral shear across Owens Valley since 83 Ma. On the order of 10 km of this displacement accumulated in the modern dextral transtension regime, and the remainder probably accumulated in Late Cretaceous–early Tertiary time. The modern tectonic regime therefore appears to reﬂect reactivation of a long-lived throughgoing crustal boundary. QUATERNARY TECTONISM OF THE NORTHWESTERN COSO RANGE The Coso Range is a tectonically and volcanically active region along the southeastern margin of the Sierra Nevada (“Sierran”) microplate, which moves ~13 mm/yr northwest with respect to stable North America (Argus and Gordon, 1991, 2001; Dixon et al., 1995, 2000). Northwest motion of the Sierran microplate is accommodated by distributed strike-slip and normal faulting in a 100-km-wide zone of active deformation

bordering the eastern Sierra Nevada (Fig. 1; Dokka and Travis, 1990; Unruh et al., 2003). Active crustal extension in the Coso Range primarily is driven by a releasing transfer of dextral motion from the Airport Lake fault to the Owens Valley fault, two major right-lateral strike-slip faults that form the eastern tectonic boundaries of the Sierran microplate south and north, respectively, of the Coso Range (Fig. 1; Monastero et al., 2005; Unruh et al., 2002). In detail, the Airport Lake fault zone splits into several branches at the southern end of the step-over region (Fig. 9). An eastern branch, consisting of N- to NNE-striking normal faults in northeastern Indian Wells Valley, extends northward across eastern Coso Basin and into the southern end of Wild Horse Mesa, where it continues northward and forms a dramatic series of scarps in Pliocene volcanic ﬂows. The faults with the largest scarps in Wild Horse Mesa exhibit a distinct left-stepping pattern (Fig. 9). A central branch, consisting of a zone of short NNE-striking, left-stepping surface traces, crosses the White Hills anticline south of Airport Lake playa and becomes the Coso Wash fault, which is characterized by a single trace along the southeastern ﬂank of the Coso Range. A western branch of the Airport Lake fault zone crosses the southern Coso Range and joins the Little Lake fault zone in southern Rose Valley (Fig. 9). Slip Transfer across the Central Coso Range Based on the geomorphic expression of faults in late Quaternary deposits, the bulk of Holocene deformation in the Coso Range appears to be associated with the central branch of the Airport Lake fault zone (Fig. 9). The most signiﬁcant structure in this branch is the Coso Wash fault, which consists of a series of NNE-striking normal faults that dip both to the SSE and the WNW. This fault zone extends from the White Hills anticline northward to Haiwee Spring in northern Coso Wash (Fig. 9) and is interpreted to be the principal locus for transferring active dextral shear through the Coso Range (Unruh et al., 2008). The Coso Wash fault zone can be traced ~9 km north of Airport Lake playa as a single SE-dipping trace that has excellent geomorphic expression as scarps in Holocene alluvial fan deposits. The fault along this reach consists of a series of alternating short NNE- and NW-striking reaches. At the southern margin of the Coso geothermal ﬁeld (GF; Fig. 9), the fault splays out into a series of WNW-dipping traces that step northwest (left) into the bedrock of the Coso Range and dip toward the main geothermal production zones. The WNW-dipping fault segments are geomorphically well-expressed by NW-facing scarps in bedrock and alluvium, and the faults locally pond alluvium in their downdropped hanging wall blocks upstream of the scarps. At the latitude of the geothermal ﬁeld, the E-dipping Coso Wash fault and the WNW-dipping normal faults collectively bound a prominent NNE-trending basement ridge that separates Coso Wash from the main production area of the Coso geothermal ﬁeld (Walker and Whitmarsh, 1998). The basement ridge is at least 10 km long (Fig. 9), locally exhibits up to 550 m of relief,

Active tectonics of the eastern California shear zone Cenozoic

1 kilometer

Mesozoic

N

Independence dike granite diorite/granodiorite

80 47 36 32

69

81

Figure 8. Simpliﬁed geologic map of the Little Lake area (modiﬁed from Bartley et al., 2008). Distribution of granite and diorite-granodiorite units east of the Owens Valley fault zone from Whitmarsh (1998). Stereoplot shows foliation (great circles) and lineation (dots) in shear zones east of the fault zone.

granodioritic gneiss (locally migmatite)

60 60

40

57

breccia zone

N

Owens Valley fault zone 25

22

shear zone (dextral reverse) 81

gneissic foliation

36 32

mylonite folation/lineation

118°0'W

118°30'W

A

B

UCF LCF

MF HS

WHM

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FZ

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US

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ARG

WHM

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GF

E

RA

Y

LE

L VA

SIER

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Lit tle

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ke

La ult fa

COSO BASIN

CB EB

WB

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INDIAN WELLS VALLEY ke t La e por Air lt zon fau

Inyokern

Ridgecrest

0

10 km

Figure 9. Northward branching of the Holocene-active Airport Lake fault zone in northern Indian Wells Valley, Rose Valley, the Coso Range, and Wild Horse Mesa. AL—Airport Lake playa; BR— basement ridge; CB—Central branch; CWF—Coso Wash fault; EB—Eastern branch; GF—geothermal ﬁeld; HS— Haiwee Spring; LCF—Lower Cactus Flat; MF—McCloud Flat; UCF—Upper Centennial Flat; WB—Western branch; WHA—White Hills anticline; WHM—Wild Horse Mesa; WHMFZ— Wild Horse Mesa fault zone. Faults with especially prominent scarps in Wild Horse Mesa are highlighted in bold. Late Quaternary faults modiﬁed from Dufﬁeld and Bacon (1981) and Whitmarsh (1998), with additional original mapping. A and B indicate two faults that display evidence for late Quaternary dextral offset (please see text for discussion).

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and is best expressed between the geothermal ﬁeld and Haiwee Spring. The basement ridge is essentially a horst block and may be analogous to the “central ridge” that Dooley et al. (2004) observed in scaled analog models of transtensional releasing stepovers that include a ductile substratum beneath a simulated brittle upper crust (quartz sand). North of the geothermal ﬁeld, the Coso Wash fault dips consistently ESE and can be traced as a series of E-facing scarps in Holocene alluvium northward to the area around Haiwee Springs, where it loses its surface expression (Fig. 9). North of Haiwee Springs, Coso Wash terminates as a Quaternary basin and narrows to a steep canyon cut in Cretaceous bedrock and Pliocene basalts of Wild House Mesa. Analysis of stereo aerial photography of this segment of the fault indicates E-facing bedrock scarps and possibly fault-related E-facing bedrock slopes. These features probably represent Quaternary faulting, as recognized earlier by Walker and Whitmarsh (1998). The step-faulted terrain associated with the eastern branch of the Airport Lake fault and the Coso Wash fault appear to merge at this location to form a rhombic array of faults at the southern end of Upper Centennial Flat (Fig. 9), north of the termination of the basement ridge. At the latitude of Haiwee Spring, the locus of active deformation steps west from Coso Wash to a series of NNE-striking, leftstepping normal faults that bound the western margins of Quaternary basins in the northwestern Coso Range such as McCloud Flat and Lower Cactus Flat (Fig. 10). The geomorphic expression and relative activity of these structures appear to increase northward as slip dies out on the Coso Wash fault and basement ridge to the east. The left step in the locus of deformation across the Coso Range is associated with an elongated, NW-trending zone of low P-wave and S-wave velocities in the depth range of ~5–12 km (Hauksson and Unruh, 2007; Reasenberg et al., 1980; Wilson et al., 2003). The base of seismicity is distinctly elevated above the low velocity zone (Monastero and Unruh, 2002), suggesting that hot ﬂuids (brines and/or magma) are present below ~5 km (Hauksson and Unruh, 2007; Wilson et al., 2003). At least two faults that display evidence for late Quaternary dextral offset can be traced north from Lower Cactus Flat into the northwestern Coso Range piedmont (faults A and B; Figs. 9 and 10). The easternmost of the two faults (A) transfers slip in a restraining stepover across NW-plunging, basementinvolved anticlines to the Red Ridge fault zone, a zone of short NNE-striking, en echelon normal faults that extends from Red Ridge northward to the margin of Owens Lake basin (Fig. 10). The westernmost of the two faults (B) is part of a zone of short, discontinuous NW-striking fault segments that trend toward the southern end of the Owens Valley fault (Fig. 10). Slemmons et al. (2008) documented dextral offset of Holocene beach ridges of Owens Lake along this fault trend, which they attribute to surface rupture during the 1872 Owens Valley earthquake. A Pleistocene pediment that fringes the northern Coso Range is deformed by WNW-trending folds in the triangular region between faults A and B (Fig. 10). The coeval NNE-SSW shortening and

WNW-ESE extension of the pediment surface in the triangular region between fault trends A and B is consistent with northwest dextral shear passing through the northwestern Coso Range. The Red Ridge fault zone, which is mapped in detail by Slemmons et al. (2008), may transfer slip to the Central Valley fault zone, a NW-striking fault in Owens Lake basin interpreted by Neponset Geophysical Corporation and Aquila Geosciences (1997) from analysis of seismic reﬂection data (Fig. 10). The “Central block” bounded by this structure and the Owens Valley fault to the west has subsided episodically during the Quaternary (Neponset Geophysical Corporation and Aquila Geosciences, 1997), possibly during large earthquakes. The NW-trending folds in the pediment surface may be the surface expression of reverse slip on blind, WNW-striking splays of the southern Owens Valley fault, also mapped by Neponset Geophysical Corporation and Aquila Geosciences (1997) from interpretation of reﬂection data (Fig. 10). Southern Owens Valley Fault, Northwest Coso Range The ﬁeld trip stop along CA-190 (Stop 7; Fig. 1) affords an excellent opportunity to look at, stand on, and walk around evidence of surface rupture during the 1872 Owens Valley earthquake. Figure 11 is a slightly oblique aerial view of the ﬁeld trip stop, looking toward the southeast. The highway parallels an abandoned shoreline of Owens Lake, and a series of gravelly late Holocene beach ridges are present on the north side of the highway. The photo shows a series of NW-trending lineaments south of the highway that project to the broad bend in the road. These features are the surface expression of the zone of discontinuous fault segments that can be traced northwest of fault A in Figure 10, and they trend toward the southern end of the Owens Valley fault zone mapped by Neponset Geophysical Corporation and Aquila Geosciences (1997) in southeastern Owens Lake basin (Fig. 10). D.B. Slemmons and his colleagues discovered that the Holocene beach ridges north of the highway are offset in a right-lateral sense along the trend of the lineaments. Slemmons et al. (2008) argue that the offset beach ridges represent surface rupture during the 1872 earthquake on the Owens Valley fault zone. If this is correct, then the 1872 rupture extended at least to the southern end of Owens Lake basin and probably south into the Coso Range piedmont (Slemmons et al., 2008). This stop is a good location to observe other tectonic-geomorphic evidence that northwest dextral shear extends from the southern end of the Owens Valley fault into the northwest Coso Range. There is a large hill ~3 km south of CA-190, on trend with the lineaments in Figure 11, where the associated fault segment terminates or makes a poorly expressed left step (Fig. 10). The hill is cut by a series of left-stepping, W-facing scarps (Fig. 12). About 3–4 km south of CA-190, there is a prominent N-facing wave-cut scarp at ~1160 m (3800 ft) elevation bordering the northern Coso Range piedmont. The shoreline of Owens Lake last reached the elevation of this escarpment ca. 24 ka during a Tioga-age pluvial highstand (Bacon et al., 2006). The shoreline

Active tectonics of the eastern California shear zone

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Figure 10. Hillshade map showing extent of a Quaternary pediment surface that fringes the northern and northwestern Coso Range. The pediment is crossed by two zones of faulting that can be traced northward from Lower Cactus ﬂat (i.e., faults A and B). The easternmost of the fault zones (A) strikes N to NNE and joins the Red Ridge fault zone in the northern Coso Range piedmont, which in turn may merge with the “Central Valley fault zone” in southern Owens Lake basin (faults in Owens Lake basin from Neponset Geophysical Corporation and Aquila Geosciences, 1997). Fault B to the west is part of a discontinuous NW-striking zone that can be traced to the southern end of the Owens Valley fault and that experienced surface rupture during the 1872 Owens Valley earthquake (Slemmons et al., 2008). The pediment surface is folded into a series of anticlines about WNW-trending axes west of the Red Ridge fault zone.

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Shor

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Figure 11. Slightly oblique aerial view to the southeast of lineaments (arrows) associated with the southern Owens Valley fault zone (distance between arrows is ~1.1 km; see Fig. 10 for location). The ﬁeld trip stop is along CA-190 at the curve in the road (Stop 7, Fig. 1). Beach ridges on the north side of the highway associated with an older Holocene shoreline of Owens Lake are offset in a right-lateral sense by the Owens Valley fault (the displacement is not visible at this scale). Slemmons et al. (2008) interpret that the displacement occurred during the 1872 Owens Valley earthquake, indicating that coseismic rupture on the Owens Valley fault extended at least as far south as the Coso Range piedmont.

Figure 12. Oblique aerial view to the east of a hill located ~3 km southeast of ﬁeld trip Stop 7 on CA-190. The west ﬂank of the hill is cut by a series of left-stepping splays of the fault shown in Figure 11. The fault splays are expressed as W-facing scarps (shadowed). Distance between the arrows is ~600 m.

Active tectonics of the eastern California shear zone scarp is visibly warped by a series of low-amplitude folds west of the Red Ridge fault zone (Fig. 10); the fold deformation is best observed from CA-190 in low-angle, early morning light when the escarpment is shadowed (Fig. 13). Although not readily visible from the ﬁeld trip stop, the Red Ridge fault zone is expressed as a series of horst and graben in a middle (?) to late Pleistocene pediment surface south of the 1160 m shoreline escarpment (Figs. 13 and 14). Summary Northwest-directed dextral shear along the southeastern margin of the Sierran microplate is transferred from the Airport Lake fault to the Owens Valley fault across a discontinuous series of active structures in the central and northwestern Coso Range. Holocene surface faults in this region locally are separated by en echelon steps, and by short restraining stepovers characterized by uplift and folding. Observations by Slemmons et al. (2008) indicate that surface rupture during the 1872 earthquake on the Owens Valley fault extended at least as far south as the late Holocene shoreline of Owens Lake, and possibly southward into the Coso Range piedmont. PLEISTOCENE DEFORMATION IN NORTHERN OWENS VALLEY Several of the broad, outstanding issues concerning the pace and tempo of active deformation throughout the eastern California shear zone–Walker Lane region exist in microcosm within the

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Owens Valley (Fig. 1). Chief among these is how one interprets differences between modern strain ﬁelds, typically measured with space geodesy (e.g., Dixon et al., 2000), and fault slip rates measured over millennia (e.g., Beanland and Clark, 1994). It has long been recognized that geodetic strain rates across the Owens Valley fault are rapid (Savage and Lisowski, 1980, 1995), and simple models of elastic strain accumulation require ~6–7 mm/yr of shear at depth to explain these observations (Gan et al., 2000). These results are in conﬂict with paleoseismic estimates of late Pleistocene to Holocene slip along the Owens Valley fault zone of 1–3 mm/yr (Bacon and Pezzopane, 2007; Beanland and Clark, 1994; Bierman et al., 1995; Lee et al., 2001b; Lubetkin and Clark, 1988). Recent studies explain this discrepancy as a consequence of transient post-seismic velocities induced by the 1872 Owens Valley earthquake (Dixon et al., 2003; Malservisi et al., 2001). However, long-term slip rates derived from paleoseismic data are subject to numerous epistemic uncertainties regarding the ratio of vertical to horizontal slip, uniform or non-uniform recurrence intervals, and the characteristic size of rupture events. Thus, such data are most useful when combined with geologic estimates of slip rate that average displacement over multiple seismic cycles. Active deformation across Owens Valley has also engendered considerable debate regarding longer-timescale variability of fault slip. At issue here is the question of whether coeval normal faulting along the Sierra Nevada range front and strikeslip faulting along the Owens Valley fault reﬂect changes in the regional stress ﬁeld (Bellier and Zoback, 1995) or simply slip partitioning along a transtensional fault system (Wesnousky and Jones, 1994). Although sites where multiple, well-dated markers

INYO MOUNTAINS COSO RANGE

Figure 13. Oblique aerial view to the east of the northern Coso Range piedmont and southern Owens Valley. The darker, more eroded surface to the south (right) is a pediment surface cut across tilted strata of the Pliocene Coso Formation. The pediment terminates to the north against a wave-cut escarpment associated with one or more Pleistocene high stands of pluvial Owens Lake. The pediment and wave-cut scarp are noticeably folded about WNW-trending axes into a series of broad, low amplitude anticlines (distance between anticline axes is ~3.2 km). As demonstrated by the varying width of the shadowed escarpment, the scarp is higher across the axes of the anticlines and lower through the syncline. See Figure 10 for location of folds relative to faults cutting the pediment surface.

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LATE

Figure 14. Oblique aerial view to the southeast of the northern Coso Range pediment surface broken up in a series of horst and graben by the Red Ridge fault zone (see Fig. 10 for location). The distance along the labeled shoreline scarp is ~1 km. PLEIS TOCE NE S HORE LINE SCAR P

are displaced across a single fault are rare, several exist within the northern part of Owens Valley. Along the Fish Springs fault, a normal fault that splays from the Owens Valley fault south of the town of Big Pine (Fig. 15), throw appears to have been steady at rates of 0.2–0.3 mm/yr over the past ~300 k.y. (Martel et al., 1987; Zehfuss et al., 2001). Likewise, normal faulting along the Sierra Nevada frontal fault appears to have been steady (throw rates of 0.2–0.3 mm/yr) when averaged over the past ~120 k.y. but may have accelerated in the late Holocene (Le et al., 2007). On longer timescales, Gillespie (1991) argued that extension along the Sierra Nevada range front underwent a period of rapid slip in the middle Pleistocene, near the time of eruption of the Bishop Tuff (ca. 760 ka; Sarna-Wojcicki et al., 2000). Similarly, coordinated variations in oblique slip rate during the middle and late Pleistocene are argued to have occurred on the White Mountains fault zone and Fish Lake Valley faults (Kirby et al., 2006; Reheis and Sawyer, 1997). Resolution of the timescales over which such variations may have occurred, and the processes driving such behavior, requires increasingly precise chronology of fault slip over multiple temporal intervals. The ﬁnal insight that active deformation within Owens Valley can provide with regard to the eastern California shear zone as a whole is the question of the role played by distributed arrays of faults in accommodating active deformation. Although active extension across the southern Owens Valley appears to be concentrated on the Sierra Nevada frontal fault (Le et al., 2007), numerous subsidiary faults occur throughout the northern valley (Fig. 15), suggesting the possibility that extension rates vary from south to north. Recent work in northern Owens Valley addresses aspects of each of these three issues. The key to surmounting uncertainties

in paleoseismic estimates of slip rate, and to assessing spatial and temporal variations in fault slip, is to quantify fault slip over millennial timescales, long enough to average multiple seismic cycles yet short enough to capture potential variations in fault slip. Direct dating of landscape features, enabled by cosmogenic isotopic methods (Gosse and Phillips, 2001), affords this opportunity (e.g., Frankel et al., 2007a; Kirby et al., 2006; Oskin et al., 2007). In the following, new geologic estimates of fault slip along the primary strand of the northern Owens Valley fault are presented using displaced lava ﬂows along the eastern ﬂank of the Crater Mountain volcanic complex (Kirby et al., 2008). Preliminary results of ongoing efforts to develop a budget of late Quaternary extension across the valley are also presented, along with a brief discussion of the regional implications of this new work. Slip Rate along the Northern Owens Valley Fault For most of its length, the trace of the 1872 rupture along the Owens Valley fault runs near or within the ﬂoodplain of the Owens River, and long-term markers of fault displacement are rare (Beanland and Clark, 1994). South of the town of Big Pine (Stop 10; Figs. 1 and 15), the fault displaces basaltic lava ﬂows along the eastern ﬂank of Crater Mountain (Beanland and Clark, 1994), one of the largest vent complexes in the Big Pine volcanic ﬁeld. Although the similarity in composition between ﬂows makes correlation of individual ﬂows across the fault difﬁcult, an apparent right-lateral separation of the contact between the ﬂow complex and alluvial fans is present at the northeastern corner of the cone (Fig. 16). The fault geometry deﬁnes a small releasing step at this site, and the pull-apart is ﬁlled with ﬁne-grained, young alluvium. Potential uncertainties regarding the degree of burial and

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Figure 15. Simpliﬁed tectonic map of the Owens Valley region, near the town of Big Pine (star). Map modiﬁed from Bateman (1965). Faults shown as solid lines are known to displace late Quaternary surﬁcial deposits, whereas those shown as dashed lines are inferred to have Quaternary slip. Box shows the location of Figure 16. CM—Crater Mountain; FSF—Fish Springs fault; OVF—Owens Valley fault; RMF—Red Mountain fault.

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the subsurface geometry of the ﬂow margin required assessment before this site could provide a robust constraint on fault slip. To test the hypothesis that dextral separation of the ﬂow margin observed at the surface (235 ± 15 m) is a reliable estimate of lateral slip along this segment of the Owens Valley fault, a ground-penetrating radar survey of the site was conducted (Kirby et al., 2008). The results of this survey conﬁrm the presence of a shallowly buried ﬂow margin to the west of the fault (Fig. 16), but do not reveal any basalt beneath the pull-apart itself. Rather, the ﬂow margin appears to end abruptly against alluvium. Thus, Kirby et al. (2008) concluded that the horizontal component of fault slip at this site is ~235 m. The age of the ﬂow surface was estimated using the concentrations of cosmogenic 36Cl in samples taken from well-preserved remnants of the ﬂow surface. Samples were selected to minimize the chance of burial by eolian or alluvial material and were collected from outcrops that exhibited glassy surfaces representative of minimal surface lowering during weathering. Three samples were collected locally, on the west side of the fault, and three additional samples were collected from ﬂows on the southwestern side of the vent complex, near the Red Mountain fault (Fig. 15). Because perfect surface preservation in this environment is extremely unlikely, exposure ages were modeled for a range of possible lowering rates (Kirby et al., 2008), and the resultant age ranges are taken as a best estimate of the age of the ﬂow. Ages from both sample localities overlap within uncertainty, and indicate that the ﬂow is 70 ± 14 ka (Kirby et al., 2008). When combined with the displacement of the ﬂow margin, these results imply that slip rates on the Owens Valley fault during this time period have been 3.6 ± 1.0 mm/yr. Notably, this range is consistent with, but at the high end of, previous estimates (Beanland and Clark, 1994; Lee et al., 2001b), and it is 2–3 times greater than Holocene estimates of Bacon and Pezzopane (2007). Whether these differences reﬂect spatial differences in slip between the northern and southern segments of the fault zone, or whether they represent a period of rapid slip prior to ca. 20–25 ka remains unknown (Kirby et al., 2008). Distributed Extension across Northern Owens Valley Ongoing work is focused on developing a budget for late Pleistocene extension across the northern part of Owens Valley. Here, a limited portion of this effort is discussed, focused on new estimates of the slip rate along the Red Mountain and Birch Mountain faults (Fig. 15). The Red Mountain fault is a 10-km-long normal fault that extends south from the southwestern ﬂank of Crater Mountain, subparallel to and ~2 km west of the Fish Springs fault (Stop 9; Figs. 1 and 15). It is marked along much of its trace by W-facing scarps that pond young alluvium in the hanging wall block. Despite its length and proximity to the Fish Springs fault, the Red Mountain structure has received considerably less attention, and the role of this fault in the extension budget across northern Owens Valley is essentially unknown.

At the northern end of the fault, lava ﬂows from Crater Mountain are displaced in a W-side–down sense and have been subsequently buried by young alluvial fans emanating from Birch Creek. The ﬂow itself buries an older alluvial fan surface with a moderately well-developed soil proﬁle; surface characteristics of this fan suggest that it is likely equivalent to surfaces dated 2 km to the east at ca. 135 ka (Zehfuss et al., 2001), whereas the younger fan appears continuous with surfaces dated at 13–15 ka (Zehfuss et al., 2001). Chlorine-36 ages from the footwall exposures of the ﬂow cluster at 70 ± 14 ka (Kirby et al., 2008) and are consistent with the regional chronology of Zehfuss et al. (2001). The ﬂow surface has been offset vertically by 14.5 ± 1.5 m, and two small scarps are present in the youngest late Pleistocene surface that exhibit a combined throw of 2.5 ± 0.2 m. Both of these displacements are consistent with long-term vertical displacement rates along the Red Mountain fault of 0.2–0.3 mm/yr (Greene et al., 2007). Thus, this structure appears to play as great a role in the regional deformation ﬁeld, as does the Fish Springs fault. Signiﬁcant slip is also present along the Sierra Nevada frontal fault at this latitude, as indicated by prominent scarps high on the range front (Fig. 17). The fault displaces sharp, triangular moraine crests in both the Tinemaha and Birch Creek drainages; moraines appear fresh, with minimal soil development and little to no weathering of boulder surfaces. Vertical displacements in both drainages are similar, ranging from 7 to 9 m (Fig. 17). Exposure ages derived from cosmogenic 36Cl concentrations in ﬁve samples taken from boulders atop the northern lateral moraine in Tinemaha Creek cluster between 14–15 ka (Greene et al., 2007). Thus, vertical displacement rates on this segment of the Sierra Nevada frontal fault system appear to be relatively high, at 0.5–0.7 mm/yr. These rates may reﬂect relatively recent breaching of the relay between the northern tip of the southern Sierra Nevada frontal fault system and the southern tip of the Round Valley fault. Regional Implications From a regional perspective, these results contribute to an emerging picture of spatial variations in slip rate along the eastern California shear zone. Recent estimates of late Pleistocene slip rates along the Death Valley–Fish Lake Valley fault system suggest that slip rates decrease northward, from 4 to 5 mm/yr along the central section (Frankel et al., 2007a) to 2–3 mm/yr along the northern sections (Frankel et al., 2007b). In a similar fashion, right-lateral slip rates in the Owens Valley appear to decrease northward, from 3 to 4 mm/yr along the Owens Valley fault to ~0.5 mm/yr along the White Mountain fault zone (Kirby et al., 2006). These apparently systematic patterns suggest that simple approaches to reconciling geologic slip and geodetic strain by summing geologic slip along a two-dimensional transect are likely to be misleading. Although strain accumulation and release may reconcile across a given transect (e.g., Frankel et al., 2007a), the spatial distribution of strain in this young, evolving fault system appears be quite variable along strike.

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Figure 17. Topographic surveys of displaced moraine crests along the Birch Mountain segment of the Sierra Nevada frontal fault zone. Surveys were conducted using post-processed differential global positioning system (sub-cm precision). Cosmogenic 36Cl ages from boulders on the Tinemaha Creek moraine indicate these deposits are 13–15 ka (Greene et al., 2007).

Distance Along Survey (m)

Summary New estimates of late Quaternary deformation rates in the vicinity of Big Pine, California suggest that (1) slip rates along the northern Owens Valley fault zone over the past 55–85 k.y. are 3.6 ± 1.0 mm/yr, signiﬁcantly higher than recent estimates along the southern portion of the fault zone, but similar to geodetic data, and (2) that extension rates associated with distributed fault arrays in the northern Owens Valley also appear to be higher than recent estimates in the southern Owens Valley. These results highlight spatial variations in late Quaternary slip rates throughout the eastern California shear zone and suggest caution in comparison of geologic slip rates with geodetic data along 2-D transects. SPATIAL VARIATIONS IN LATE PLEISTOCENE SLIP RATE ALONG THE DEATH VALLEY–FISH LAKE VALLEY FAULT ZONE The Death Valley–Fish Lake Valley fault zone is the largest and most continuous fault system in the eastern California shear zone, extending some 300 km northward from its intersection with the Garlock fault (Fig. 1). Both geologic and geodetic observations suggest that the Death Valley–Fish Lake Valley fault zone accommodates the majority of slip in the northern eastern California shear zone. Speciﬁcally, several space-based geodetic surveys show that, over the past ~15 yr, the Death Valley–Fish Lake Valley fault zone has been taking up 3–8 mm/yr of the measured 9.3 ± 0.2 mm/yr of Paciﬁc–North America plate motion in the northern eastern California shear zone and Walker Lane (Bennett et al., 1997, 2003; Dixon et al., 1995, 2000, 2003; Humphreys and Weldon, 1994; McClusky et al., 2001; Savage et al., 1990; Wernicke et al., 2004). Along the northern part of the Death Valley–Fish

Lake Valley fault zone in Fish Lake Valley, modeling of geodetic data shows the fault system is storing strain at 4–10 mm/yr and that the White Mountains fault zone, the other major strike-slip structure at this latitude, stores strain at 1–5 mm/yr (Dixon et al., 1995, 2000). The geodetic data, therefore, suggest that at latitude ~37.5°N, essentially all plate boundary strain in the eastern California shear zone is accommodated on these two structures. Although numerous geodetic data are available from the region, only a few ﬁeld-based studies have attempted to measure intermediate- and long-term (1,000–1,000,000 yr) geologic slip rates on the Death Valley–Fish Lake Valley fault zone (Brogan et al., 1991; Frankel, 2007; Frankel et al., 2007a, 2007b; Klinger, 2001; Reheis and Sawyer, 1997). A lack of geochronologic constraints on the age of offset alluvial landforms has thus made it difﬁcult to compare rates of deformation over multiple time scales along this part of the plate boundary. Previous estimates of the late Pleistocene slip rate for the Death Valley–Fish Lake Valley fault zone in Fish Lake Valley range from 1 to 9 mm/yr (Reheis and Dixon, 1996; Reheis and Sawyer, 1997). However, recent work combining high-resolution ALSM digital topographic data with cosmogenic nuclide geochronology to investigate fault offsets and slip rates has helped reﬁne the long-term slip rates in this region (Frankel, 2007; Frankel et al., 2007a, 2007b). These new results reveal spatial variations in the slip rate along the Death Valley–Fish Lake Valley fault zone and have important implications for eastern California shear zone and Paciﬁc–North America plate boundary kinematics. Faulting in Fish Lake Valley The Death Valley–Fish Lake Valley fault zone bounds the east side of the White Mountains in Fish Lake Valley (Fig. 1).

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This region makes up the northern 80 km of the Death Valley–Fish Lake Valley fault zone. Estimates of total dextral displacement along the northern Death Valley–Fish Lake Valley fault zone are thought to range from ~50 to 80 km since Cambrian to Middle Jurassic time (McKee, 1968; Stewart, 1967). More recent right lateral–oblique fault activity in Fish Lake Valley is characterized by numerous deformed geomorphic features, including fault scarps, displaced alluvial fans, offset drainage channels, shutter-ridges, and sag ponds (e.g., Brogan et al., 1991). An ALSM survey was recently conducted along the Death Valley–Fish Lake Valley fault zone in northern Death Valley and Fish Lake Valley to precisely map displacement on late Pleistocene and Holocene fault-related landforms (please see Frankel, 2007, and Frankel et al., 2007b for details about ALSM data collection and processing parameters). Results from two of the survey locations, the offset Furnace Creek and Indian Creek alluvial fans in Fish Lake Valley, are discussed below.

Furnace Creek Offset The Furnace Creek alluvial fan in central Fish Lake Valley is located along the Oasis section of the Death Valley–Fish Lake Valley fault zone (Stop 11, Fig. 1; Brogan et al., 1991; Reheis and Sawyer, 1997). At this location, alluvial fans are offset along two parallel strands of the fault (Fig. 18). Previous work established a number of possible late Pleistocene channel displacements at Furnace Creek, ranging from 111 to >550 m (Brogan et al., 1991; Reheis et al., 1995; Reheis and Sawyer, 1997). ALSM data were used to determine more precise offset measurements. These high-resolution topographic data helped reveal subtle topographic features with hillshade, topographic, and slope aspect maps, topographic proﬁles, and thalweg positions to reconstruct a prominent drainage channel and the overall fan morphology (Fig. 18). Based on these data, Frankel et al. (2007b) revised the late Pleistocene strike-parallel displacement for the Furnace Creek alluvial fan to 290 ± 20 m (Fig. 18).

118.01°W

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Figure 18. Maps of Furnace Creek alluvial fan in Fish Lake Valley (Stop 11, Fig. 1; after Frankel et al., 2007b). (A) Hillshaded 1-mresolution ALSM digital elevation model of the Furnace Creek alluvial fan. Star indicates parking area for ﬁeld trip. (B) Geologic map (modiﬁed from Reheis et al., 1995) of the Furnace Creek alluvial fan draped over the hillshaded image from A. Qﬁo—Older alluvium of Indian Creek (late Pleistocene); Qft—Alluvium of Trail Canyon (middle Pleistocene); Qfm—Alluvium of McAfee Creek (middle Pleistocene). (C) Furnace Creek alluvial fan retro-deformed 290 ± 20 m to its prefaulting position based on the high-resolution ALSM digital elevation data. Hatched pattern on northwest section of the offset fan indicates a surface of similar age, but set into the Qﬁo unit. Combining the late Pleistocene displacement history with the 94 ± 11 ka cosmogenic 10 Be model age from the offset Qﬁo surface yields a slip rate of 3.1 ± 0.4 mm/yr (Frankel et al., 2007b).

Active tectonics of the eastern California shear zone Indian Creek Offset The Indian Creek alluvial fan is located at the north end of Fish Lake Valley, near the northern termination of the Death Valley–Fish Lake Valley fault zone (Stop 12, Fig. 1). The fan is part of the Chiatovich Creek section of the Death Valley–Fish Lake Valley fault zone (Brogan et al., 1991; Reheis and Sawyer, 1997). Here, the fault zone splays into numerous normal faults, and the strike-slip component of deformation is localized along a single strand that displaces both late Pleistocene and Holocene alluvium (Fig. 19). Reheis et al. (1993) and Reheis and Sawyer (1997) estimated 83–165 m of late Pleistocene rightlateral deformation at this location based on a single offset debris ﬂow channel. Hillshade, slope aspect, and topographic maps and channel thalwegs derived from ALSM data were used to revise the late Pleistocene displacement at Indian Creek to 178 ± 20 m by retrodeforming at least four, and possibly six, offset channels incised through the alluvial fan (Fig. 19; Frankel et al., 2007b).

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Alluvial Fan Ages Ages of the offset alluvial fans in Fish Lake Valley were previously estimated on the basis of soil development and surface morphology (e.g., Reheis et al., 1993, 1995; Reheis and Sawyer, 1997). Both the Furnace Creek and Indian Creek fans have similar soil and morphologic characteristics. The offset late Pleistocene fans have well-developed soils with a 5–10-cm-thick silty vesicular A horizon and an argillic B horizon with moderate clayﬁlms and stage II to III carbonate development (Reheis and Sawyer, 1997). The fan surfaces are characterized by a moderately to well-developed desert pavement, moderate to dark desert varnish coatings on clasts ranging in size from pebbles to boulders, and subdued to moderately incised channels (Fig. 20). Ages of the offset alluvial fans at Furnace Creek and Indian Creek were quantiﬁed by cosmogenic nuclide 10Be geochronology (Frankel et al., 2007b; Gosse and Phillips, 2001). Geochronology samples were collected from the top 2–5 cm of large

B

Figure 19. Maps of the Indian Creek alluvial fan in Fish Lake Valley (Stop 12, Fig. 1; after Frankel et al., 2007b). (A) Hillshaded 1 m resolution ALSM digital elevation model of the Indian Creek alluvial fan. Star indicates parking area for ﬁeld trip. (B) Geologic map (modiﬁed from Reheis et al., 1993) of the Indian Creek alluvial fan and surrounding areas draped over the hillshaded image in A. Qﬁy—Younger alluvium of Indian Creek (late Pleistocene); Qﬂ—Alluvium of Leidy Creek (early Holocene and late Pleistocene); Qls—Holocene to late Pleistocene landslide deposits. (C) Indian Creek alluvial fan retrodeformed 178 ± 20 m to its pre-faulting position based on ALSM data. A cosmogenic 10Be model age for the Qﬁy surface of 71 ± 8 ka yields a slip rate of 2.5 ± 0.4 mm/yr when combined with the late Pleistocene offset measurement (Frankel et al., 2007b).

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Figure 20. Photograph looking west across the Furnace Creek alluvial fan. The White Mountains make up the distant skyline ridge. Boulder in the foreground is representative of locations from which samples were collected on the Furnace Creek and Indian Creek alluvial fans. Note the high degree of varnish development on the boulder surface. Boulders averaged ~115 × 90 × 85 cm at Furnace Creek and ~65 × 55 × 50 cm at Indian Creek.

boulders on the stable parts of offset fan surfaces mapped as Qﬁ by Reheis et al. (1993, 1995; Fig. 20) and analyzed by accelerator mass spectrometry at Lawrence Livermore National Laboratory (please see Frankel, 2007, and Frankel et al., 2007b, for geochronology methods and data). Furnace Creek Fan Age Eight cosmogenic 10Be samples from the Qﬁo surface at Furnace Creek (Fig. 18; unit Qﬁ of Reheis et al., 1995) range in age from 79 ± 8 ka to 112 ± 8 ka (Frankel et al., 2007b). The relatively tight cluster of dates suggests that the Qﬁo surface has remained stable, with samples having been exposed to cosmic rays in their current conﬁguration, since deposition. The age of the fan is taken to be the mean and standard deviation of the eight samples, which yields a date of 94 ± 11 ka (Frankel et al., 2007b). This age falls within the broad age range of 50–130 ka estimated for the Furnace Creek fan on the basis of soil development and surface morphology by Reheis and Sawyer (1997). Indian Creek Fan Age Eight cosmogenic 10Be samples were also collected from the Qﬁy surface (Fig. 19; unit Qﬁ of Reheis et al., 1993) at Indian Creek. These samples range in age from 59 ± 4 ka to 81 ± 6 ka (Frankel et al., 2007b). The samples from Indian Creek are somewhat younger than those from Furnace Creek, yet they display a similarly tight cluster, which is also interpreted to indicate the stability of the Qﬁy surface since deposition. The mean and standard deviation of the eight 10Be dates yields an age of 71 ± 8 ka for the Qﬁy surface at Indian Creek, which is in agreement with the previously estimated range of 50–130 ka (Frankel et al., 2007b; Reheis and Sawyer, 1997).

The cosmogenic 10Be surface exposure dates from the Furnace Creek and Indian Creek fans are interpreted as maximum ages for the deposits because the channels used as piercing points must have formed at some undetermined time following fan deposition. The slip rates reported here should therefore be interpreted as minima, though it appears that most, if not all, of the rightlateral deformation accommodated on faults has been accounted for. The pervasive normal faulting observed on the eastern White Mountains piedmont is not considered in these rates. Combining the displacement of 290 ± 20 m with the age of 94 ± 11 ka for the offset Qﬁo surface at Furnace Creek yields a late Pleistocene slip rate for the Oasis section of the Death Valley–Fish Lake Valley fault zone of 3.1 ± 0.4 mm/yr. Previous slip rate estimates at this location ranged from 1.5 to 9.3 mm/yr (Reheis and Sawyer, 1997). The 178 ± 20 m of displacement and 71 ± 8 ka age of the offset Qﬁy surface at Indian Creek results in a slightly slower slip rate of 2.5 ± 0.4 mm/yr along the northern Chiatovich Creek section of the Death Valley–Fish Lake Valley fault zone. This rate falls within the bounds of the previous slip rate estimate of 1.1–3.3 mm/yr for this site (Reheis and Sawyer, 1997). Spatial Variations in Slip Rate and Northern Eastern California Shear Zone Strain Distribution Previous studies in the northern eastern California shear zone suggest that the down-to-the-NW faults located between the major strike-slip faults of the region transfer slip from the Owens Valley and Panamint Valley–Hunter Mountain–Saline Valley faults to the northern part of the Death Valley–Fish Lake Valley fault zone (Fig. 1; Dixon et al., 1995; Lee et al., 2001a; Reheis and Dixon, 1996). However, recent results from three slip rate sites along the Death Valley–Fish Lake Valley fault zone show that this may not be the case (Frankel et al., 2007a, 2007b). The late Pleistocene slip rate along the Death Valley–Fish Lake Valley fault zone in northern Death Valley is ~4.5 mm/yr (Frankel et al., 2007a). The rates determined at the offset Furnace Creek and Indian Creek alluvial fans show that this rate decreases to ~2.5–3 mm/yr on the northern part of the Death Valley–Fish Lake Valley fault zone in Fish Lake Valley (Frankel et al., 2007b). The late Pleistocene slip rate on the White Mountains fault zone is 0.3–0.4 mm/yr (Kirby et al., 2006). Taken together, the late Pleistocene slip rates on the two major faults at latitude ~37.5°N are less than half the 9.3 ± 0.2 mm/yr region-wide rate of dextral shear determined from geodetic data (Bennett et al., 2003). This result suggests either that deformation at the latitude of Fish Lake Valley is accommodated on structures other than the White Mountains fault zone and Death Valley–Fish Lake Valley fault zone or that a strain transient exists in the northern eastern California shear zone, similar to that proposed for the Mojave Desert (Oskin and Iriondo, 2004; Oskin et al., 2006, 2007). Strain rates appear to have remained constant in the northern eastern California shear zone over the past ~70 k.y. at the

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latitude of northern Death Valley (Frankel et al., 2007a). If true, this implies that strain must be accommodated off the White Mountains and Death Valley–Fish Lake Valley fault zones, either to the west through Long Valley caldera (e.g., Kirby et al., 2006) or to the east from the Emigrant Peak fault zone through the Silver Peak–Lone Mountain extensional complex (Fig. 1; Oldow et al., 1994). If the faults to the east are indeed accommodating this additional strain, then the eastern California shear zone–Walker Lane transition must occur farther south, in a broader, more diffuse zone than previously recognized.

FIELD GUIDE

Summary

Day 1: Mojave Desert, Summit Range, and Red Rock Canyon

The late Pleistocene slip rate on the Death Valley–Fish Lake Valley fault zone decreases from ~4.5 mm/yr in northern Death Valley to ~3 mm/yr at Furnace Creek in central Fish Lake Valley and ~2.5 mm/yr at Indian Creek in northern Fish Lake Valley. Combining slip rates from the northern Death Valley–Fish Lake Valley and White Mountains fault zones, the two major faults at latitude ~37.5°N, indicates that the late Pleistocene rate of dextral shear is less than half that determined from geodetic data. This suggests either the existence of a strain transient in the northern eastern California shear zone or that deformation is distributed across other structures in the region.

Stop 1: Lenwood Fault (Southern Offset) Directions. This stop is located at UTM 507142E, 3844558N in Stoddard Valley, east of highway 247, ~10 mi south of Barstow, California. To reach the site from Interstate 15, exit at CA247 (Barstow Road). Turn south on CA-247 and drive ~10 mi. The ﬁeld site is located on Bureau of Land Management (BLM) route OM8, which is ~1 mi south of the Slash X Ranch Café, on the left. To locate the road, look for the end of barbed wire fence and a “Call Box” sign. Turn left on to BLM route OM8 immediately after the “Call Box” sign. There are several homesteads on OM8 on the way to the site, so please drive slowly to keep the dust down and watch for oncoming trafﬁc. Take OM8 for ~4.5 mi. You will pass over a cattle guard. If you drive through a creek bed, then up a steep rocky section of the road ~4 m high, you’ve probably just passed it. Park in the ﬂat area just before this rocky slope (Fig. 2). Description. Figure 2 shows a hillshade map of this stop generated from a portion of the Lenwood fault ALSM survey (also see Fig. 21). A feature of particular interest at this location is a 100-m-wide pull-apart basin with slip partitioned onto several strands. A channel that crosses this basin appears offset ~45 m. Northwest of the pull-apart basin, fault displacement

SUMMARY The eastern California shear zone accommodates the majority of Paciﬁc–North America plate boundary motion east of the San Andreas fault. Therefore, deformation in this region is critical to our understanding of plate boundary kinematics, in addition to the behavior and evolution of fault systems. The studies presented in this guidebook are some of the most recent investigations undertaken in the region. We have highlighted new late Pleistocene slip rates in Owens Valley, Fish Lake Valley, and the Mojave Desert; discussed the timing and rates of offset along the Garlock fault; provided evidence for ~65 km of total cumulative right-lateral displacement across Owens Valley; and observed recent deformation associated with southern Owens Valley fault in the northwestern Coso Range. Together, the topics in this guidebook span the Cretaceous to late Holocene tectonic history of the eastern California shear zone. As is often the case, much of this recent research, while generating a wealth of new data on slip rates, displacement histories, and fault kinematics, has brought about even more questions regarding spatial and temporal patterns of deformation in the eastern California shear zone. The eastern California shear zone provides a unique opportunity to study the kinematics of an evolving plate boundary. Continued work in the region, focused on deﬁning rates of deformation across broad temporal (tens to millions of years) and spatial (tens to hundreds of kilometers) scales, will help ﬁll gaps in our understanding of the role the eastern California shear zone plays in accommodating Paciﬁc–North America plate boundary deformation.

This guide provides directions to ﬁeld trip stops relative to nearby towns, landmarks, and major roads. It does not provide distances between individual stops. All ﬁeld trip stop locations are keyed to the map in Figure 1. In addition, Universal Transverse Mercator (UTM) coordinates of all stops are provided in NAD83 datum, zone 11. It is preferable to have a high-clearance vehicle (and possibly 4WD, depending on road conditions) for many of the stops.

Figure 21. Low-relief scarps of the Lenwood fault (in mid-ground; Stop 1). Photograph taken looking to the northeast.

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formed a prominent SW-facing scarp cutting both Q2a and Q2b alluvial fans. Inset into the fans are younger channel-ﬁll deposits also cut by low, SW-facing scarps in several localities. One of the larger channels crossing the fault is deﬂected ~270 m. Most of this deﬂection is due to construction of a Q2b alluvial fan in front of the channel. Two secondary faults form scarps in the late Pleistocene fans east of the main fault trace.

valley. The slip rate of the Lenwood fault was estimated here from deﬂection of a pair of channels located immediately NW of the road, and from a shutter ridge forming in front of the larger channel descending the fault-line valley. These channels are inset into outcrops of conglomerate and Q2b alluvial fan surfaces. Based on the 37 ± 7 ka age of Q2b determined from the southern site, the slip rate here is 0.8 ± 0.2 mm/yr.

Stop 2: Lenwood Fault (Northern Offset) Directions. This stop is located at UTM 501464E, 3851618N. The stop is ~3.5 mi from the highway. From Stop 1, return to CA-247 and turn right to go north. Approximately 1.5 mi north of the Slash X Ranch Café, bear right onto Stoddard Valley Road. The intersection is located 0.2 mi past the power lines with towers shaped like a Π symbol. Take this road straight back with no turns until you are past the ﬁrst set of hills. Here, look for a dirt track cutting southeast, up the valley, to join the power-line road. Turn left onto the power-line road and ascend the hills. Note that it is also possible to drive the power-line road from CA-247; however, this route requires a very steep descent of the ﬁrst set of hills. Once on the power-line road, stay left at the next three forks. Park at the fourth and ﬁnal fork, located at the crest of the hills (Fig. 3). The Lenwood fault runs along the hillside, ~80 m below. Walk down the road to reach the fault. It is not recommended to drive down the hill, as the descent is dangerously steep and may require 4WD to ascend. Description. Stop 2 features an ~600 m length section where the Lenwood fault is spectacularly exposed as a set of uphill-facing scarps and shutter ridges (Figs. 3 and 22). Southward, the fault crosses a wash and runs up the northeast side of a fault-line

Stop 3: Garlock Fault in Summit Range Directions. This stop is located at UTM 445564E, 3924572N. This location is between Barstow and Ridgecrest, near the small town of Johannesburg. From Barstow, head east on CA-58 until it intersects with US-395. Go north on US-395 for ~27 mi to Trona Road. Turn east on Trona Road and continue for ~7.75 mi to the dirt road leading off to the west; park there. Description. This stop is to introduce the main units in the Summit Range. This will be the basis for correlation/comparison with rocks in the Red Rock Canyon area (Stop 4). We will walk through the stratigraphy here to examine the various units (Figs. 5 and 23). First stops will be in the lower sedimentary sequence. We will then examine the lapilli tuff and its relation to underlying and overlying units. Lastly, we will look at the various tuffs overlying the lapilli tuff. Stop 4: Garlock Fault in Red Rock Canyon Directions. This stop is located at UTM 411017E, 3913555N. Red Rock Canyon is located 25 mi north of Mojave on CA-14, near Cantil. Signs on CA-14 clearly indicate the turnoff on Abbott Road. Go west on Abbott Road to the visitor center parking lot and park. From there, walk ~0.2 mi northeast to the outcrop.

Figure 22. View looking to the northwest at fault scarps and shutter-ridges along the Lenwood fault (Stop 2). Note truck on right side of photograph for scale.

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Description. This stop will emphasize units of the Dove Spring Formation (Figs. 5 and 24). We start at the parking area east of the road and proceed up section. Our ﬁrst stops are in the sedimentary and volcanic units below the lapilli tuff. First outcrops are the conglomerates in the Td2 sequence. We will then work our way up-section into white tuffs below the thick lapilli tuff. Proceed then through the tuff looking at the possible cooling break in the middle of this massive unit. We will then look at (or hike westward to) the rocks above the tuff, including the basalt ﬂows in the higher parts of the section. Day 2: Little Lake, Cactus Flat, Coso Range, Independence Creek, and Birch Creek Stop 5: Independence Dike Swarm at Little Lake Directions. This stop is located at UTM 417369E, 3977338N, just west of Little Lake. Heading north on US-395, exit at the southernmost Little Lake Road exit. (NOTE: if approaching this stop from the north, this will be the second Little Lake Road exit; if approaching from the south, this will be the ﬁrst Little Lake Road exit.) Cross to the west side of the highway and drive 0.1 mi for a view stop. Description. This is stop 13 of Glazner et al. (2005). The Little Lake structural block (Fig. 8) exposed at this stop and in the mountains directly to the east is composed of Jurassic (165 Ma; Whitmarsh, 1998) plutonic rocks that vary in composition from diorite to leucogranite (mainly as a result of pervasive magma mingling) and lack a penetrative fabric. The heterogeneous plutonic complex is intruded by Jurassic (148 Ma; Whitmarsh, 1998) Independence dikes that vary in composition roughly as much as their wall rocks do. Because the dikes are more erosion-resistant than their wall rocks, most of the craggy outcrops on the hills to the east are dikes. A traverse of the skyline to the southeast of the stop yielded 9.1% dilation by dike intrusion (Fig. 8; Bartley et al., 2008). Individual dikes in the Little Lake block range up to >10 m thick, yet any individual dike cannot be traced for more than

Figure 23. Photograph looking southeast across the Summit Diggings volcanic ﬁeld (Stop 3). The prominent ridge in the center of the photo is a N80E-striking, SSE-dipping sequence of tuffs capped by a distinctive pink lapilli ash-ﬂow tuff. This outcrop is surrounded by a patchwork of tuffs, epiclastic rocks, and sedimentary rocks of the Bedrock Springs Formation. The photo is taken from the side of a dacite dome (dark-colored rocks in left foreground) which is part of a larger 12– 11 Ma dome-ﬂow-tuff complex. The tuffs occur in a topographically low valley that is likely related to formation of a small caldera resulting from the extrusion of the pink ash-ﬂow tuff.

200 m because the dikes are transected by numerous NW- to N-striking ductile-brittle shear zones. The complete mismatch of dikes across each shear zone suggests that displacements across individual zones may be large. Shear zones commonly are exposed in numerous prospect pits that dot the area, and consist of greenschist-facies (white mica–chlorite–albite–epidote ± biotite) phyllonite. Most of the shear zones dip steeply to moderately westward (Fig. 9), and lineation and shear-sense indicators (S-C composite foliations, asymmetric porphyroclasts, mica ﬁsh)

Figure 24. An ~125 m sequence of tuffs, epiclastic, and clastic units of the lower Dove Spring Formation, Ricardo Group, located on the north side of the Garlock Fault in the Red Rock Canyon State Park (Stop 4). The E-dipping normal fault through the sequence displaces the units ~25 m. The prominent white air fall tuffs are dated at older than 11 Ma. The capping pink lapilli ash-ﬂow tuff correlates with the pink tuff found at the Summit Range, 34 km to the east on the south side of the Garlock Fault. The red and tan clastic units are also identical in lithology to those found in the Summit Range area that are attributed to the Bedrock Springs Formation.

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indicate varying proportions of dextral and reverse displacement (Bartley et al., 2008). Shear zones decrease in abundance eastward toward the main axis of the Coso Range. Timing relations, kinematics, and spatial distribution of the shear zones suggest that they may be representative of structures that accommodated signiﬁcant offset across Owens Valley. Roadcuts along the microwave tower access road south of this stop expose rocks of the Sierran block, including granodiorite and granodiorite gneiss, variably intruded by 10–20-cm-thick maﬁc dikes, all of which are thoroughly shattered. Like plutonic rocks of the Little Lake block, the granodiorite yielded a U-Pb age near 165 Ma (Bartley et al., 2008), but geologic similarities between the Little Lake block and adjacent Sierra rocks end there. The granodiorite varies much less in composition than Little Lake plutonic rocks, contains a pervasive gneissic fabric that is locally migmatitic, and contains quartzite and calc-silicate rock xenoliths not seen in the Coso Range. Dikes in the granodiorite are uniformly maﬁc, rarely as thick as 1 m, vary widely in orientation, and account for only 1%–2% dilation. In areas where geochronology and cross-cutting relations distinguish Jurassic Independence dikes from associated Cretaceous dikes (Coleman et al., 2000; Mahan et al., 2003), all of these characteristics correspond to Cretaceous dikes. Note that rocks of the Little Lake block are exposed on the west side of US-395 in low outcrops along the west side of Little Lake Road immediately north of here (on and around hill at UTM 417979E, 3977298N). There is an almost complete geologic mismatch between the Little Lake block and the adjacent Sierra Nevada, and it is inferred that the boundary in between, which is concealed by alluvium, is a locus of major tectonic offset.

Stop 6: Coso Dike at Cactus Flat Directions. This stop is located at UTM 418900E, 4008439N, ~5 mi southeast of Olancha. From Olancha, drive south on US395 ~1.5 mi to Cactus Flat Road (opposite the ﬁre station). Turn east and drive past the pumice mine (mine is 4.2 mi from intersection of Cactus Flat Road and US-395). Two and a half miles beyond the mine, make a left turn on a narrow dirt road, and 0.4 mi beyond this, turn left at a side road at UTM 419210E, 4007018N. Continue down this road ~0.5 mi and park at UTM 419490E, 4007610N. An outcrop of the Coso dike is obvious in the hillside north-northwest of the parking area. Description. This is Stop 11 of Glazner et al. (2005). Cretaceous dikes in the Coso Range were ﬁrst described by Dufﬁeld et al. (1980) as steeply dipping, W-striking, conspicuously K-feldspar and quartz-phyric granitic dikes that are up to 10 m thick. Whitmarsh (1998) named these dikes the Coso dike swarm, mapped the dikes in signiﬁcant detail, and obtained a preliminary U-Pb zircon date of ca. 84 Ma from a thick dike in the swarm exposed near Upper Centennial Flat. One of us (Glazner) made a reconnaissance trip into the Coso Range in 1996 and noted the strong similarity between the Coso and Golden Bear dikes and proposed that the two areas might once have been contiguous. Kylander-Clark (2003) tested this correlation and found identical ages of 83.5 Ma for the two swarms, overlapping geochemistry, and similar, unique wall rocks (102 Ma leucogranite) that strongly support correlation of the two areas and ~65 km of dextral displacement. At this locality, the Coso dike (Fig. 25) is >20 m wide, strikes 281°, and dips ~80°N. The quartz monzonite porphyry contains 1–5-cm-long euhedral K-feldspar phenocrysts that locally contain Carlsbad twins. Plagioclase occurs as small (2– 4 mm) subhedral, subequant grains. Quartz occurs in euhedral

Figure 25. Westernmost outcrop of Coso dike swarm in the Coso Range (Stop 6). View is to the northwest from Cactus Flat. The dike is ~20 m thick, strikes W, and forms the prominent crags on skyline; the skyline behind is underlain by Pliocene maﬁc volcanic rocks that overlap the dike.

Active tectonics of the eastern California shear zone bipyramids that range from 2 to 8 mm in diameter. Biotite is the major maﬁc mineral, occurring as small (typically 1 mm) subhedral to euhedral grains. Looking east, the dike continues on strike, but is difﬁcult to follow from a distance. Stop 7: Northern Coso Range Piedmont Directions. This stop is located at UTM 413630E, 4019852N. From Olancha, head east on CA-190, ~3.5 mi from the intersection with US-395 (~1 mi west of the well-marked road to Dirty Sock Corporation Yard) and park on the side of the road along the broad shoulder. Description. The thing to look at, stand on, and walk around here is evidence for surface rupture during the 1872 Owens Valley earthquake. With a little searching, you will ﬁnd a series of rightlaterally displaced Holocene beach ridges on the north side of the highway that were discovered by Burt Slemmons and his colleagues ~30 years ago and that are on trend with the Owens Valley fault to the north where it heads into the playa from Bartlett Point (Fig. 11). Slemmons et al. (2008) argue that the offset beach ridges indicate that surface rupture during the 1872 earthquake extended all the way to the Coso Range piedmont, rather than dying out at Bartlett Point as suggested by Beanland and Clark (1994). Once you ﬁnd the offset beach ridges, look to the southeast to view tectonic-geomorphic evidence for dextral shear extending from the southern end of the Owens Valley fault into the northwest Coso Range. There is a large push-up ridge ~3 km south of the road where the fault zone that offsets the beach ridges terminates or makes a left step. About 3–4 km south of the road is a high, N-facing wave-cut scarp at ~1160 m (3800 ft) elevation along the northern Coso Range piedmont. This scarp was last inundated ca. 24 ka during a Tioga-age pluvial highstand of

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Owens Lake. The wave-cut scarp is visibly warped into a series of low-amplitude folds associated with a fault zone that transfers slip from Owens Valley southward through the Coso Range to the Airport Lake fault zone in Indian Wells Valley. These structures comprise the eastern tectonic margin of the Sierra Nevada microplate at this latitude. Stop 8: Golden Bear Dike Near Independence Creek Directions. This stop is located at UTM 387410E, 4068329N along the Sierra Nevada range front southeast of Independence. From the center of Independence, drive 4.5 mi west on Onion Valley Road. Turn south on Foothill Road and continue for 1.5 mi to the foot of the boulder-covered hill and park on the side of the road. Description. This is Stop 6 of Glazner et al. (2005). The Golden Bear dike crops out from the valley ﬂoor to the range crest and beyond, crossing the crest north of Forester Pass to eventually peter out in the headwaters of the Kern River (Fig. 6). Looking west from this locality, it is possible to trace the dike from the foothills up the range front to the crest of the Sierra Nevada (Fig. 26). Here, in the foothills, the dike is shattered and dismembered due to range-front faulting and landsliding. However, most of the large boulders mantling the slope are distinctive porphyritic Golden Bear dike. West of the Independence fault (Moore, 1963), the dike is intact, strikes nearly E-W, dips steeply, and is 10–15 m wide. Here we will see large ﬂoat boulders of the dike and gain an appreciation for its width the last time it is exposed passing east into Owens Valley. The Golden Bear dike is a K-feldspar quartz monzonite porphyry with euhedral zoned phenocrysts of K-feldspar that range 2–4 cm in length and commonly display Carlsbad twins.

Figure 26. Golden Bear dike on the south side of Pinyon Creek drainage, eastern ﬂank of Sierra Nevada, viewed looking west from Onion Valley Road. Dashed lines follow the outcrop trace of the dike. The prominent peak, Mount Bradley (13,289 ft, 4050 m), is carved in the 102 Ma Bullfrog leucogranite, another tie point across Owens Valley. Stop 8 is along the range front just left (south) of photo.

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Plagioclase is subhedral, subequant, and typically 1–4 mm in diameter. Quartz typically occurs as distinctive equant, euhedral, and bipyramidal crystals 2–5 mm in diameter. Maﬁc minerals include both biotite and rare hornblende. Stop 9: Red Mountain Fault at Birch Creek Directions. This stop is located at UTM 384884E, 4103959N. From Big Pine, head south on US-395 ~5 mi to North Fish Springs Road. Turn west (right) on North Fish Springs Road and continue until it begins to bend back to the east (toward US395). At the bend (a large triangular intersection), head west on Tinemaha Road. Bear to the right after 300 m, at the base of the Poverty Hills. Birch Creek Road will appear as a dirt road on the right on a very gentle right-hand turn (almost straight) parallel to the creek. This road passes through a cluster of homes, crosses Birch Creek, and then turns back to the west to parallel the creek again. Continue up the switchbacks onto the alluvial fan surface. Stay to the left (straight) at the intersection with a power-line road. The road then veers to the northwest away from Birch Creek toward the granitic Fish Springs Hills. Take the ﬁrst left, following Birch Creek Road back to the southwest (toward Birch Creek). In ~0.6 mi, the road descends at a W-facing scarp. Park just west of the scarp at UTM 384964E, 4103425N. Walk ~400 m north along the scarp to the large gully. Description. Although recent extension across the southern Owens Valley, near the town of Lone Pine, is primarily accomplished by slip on the Sierra Nevada frontal fault (Le et al., 2007), the geometry and distribution of recent faulting changes markedly in the northern part of the valley (Fig. 1). North-striking normal faults occur in distributed arrays that extend from the foot of the range across the western piedmont to the Owens Valley fault. The best known of these structures is the Fish Springs fault, a W-side-down normal fault that forms the large scarp ~2 km east

of this location. Slip on this structure appears to have been downdip and relatively steady at throw rates of 0.2–0.3 mm/yr (Martel et al., 1987; Zehfuss et al., 2001) since ca. 330 ka. The association of this structure with the surface trace of the Owens Valley fault has led most workers to consider them part of the same fault zone. However, the broad array of normal faults of similar orientation north and west of this location suggests that the Fish Springs fault may be simply one of a distributed array of normal faults that accommodate extension across the Owens Valley at this latitude. The goal of this stop is to examine evidence for the rates of displacement along a couple of these structures. This stop begins at the W-facing scarp of the Red Mountain fault where it displaces alluvial fan deposits of Birch Creek (presently deeply incised to the south of the parking spot; Fig. 27). The topographic expression of the fault is not particularly impressive (a few meters), due to the fact that fan aggradation during and shortly after the last glacial maximum (Zehfuss et al., 2001) on the downthrown hanging-wall block has largely ﬁlled the accommodation space. Just north of the road, bouldery surfaces of this fan spill over the scarp and are continuous with sites to the east, dated by Zehfuss et al. (2001) at 13–15 ka (Fig. 27). Approximately 100 m north of the Birch Creek road, two small (~1 m) scarps are present within this alluvial surface. These likely represent the most recent rupture along the Red Mountain fault, and, if representative of average slip, would imply single event throw rates of ~0.2 mm/yr. The Red Mountain fault also displaces lava ﬂows along the southwestern ﬂank of Crater Mountain. Continue northward along the fault scarp, until you reach a large gully draining through the Fish Springs Hills (~300 m). East of this gully, ﬂows are exposed capping a fan deposit with a moderately well-developed soil proﬁle. West of the gully, the ﬂow has been displaced vertically ~12–14 m and buried by younger alluvium. Cosmogenic 36Cl

Figure 27. Field photograph taken looking toward the northeast at a scarp of the Owens Valley fault zone at Birch Creek in Owens Valley (Stop 9). Ridges in the background are the Fish Springs Hills.

Active tectonics of the eastern California shear zone ages from the ﬂow surface indicate that the ﬂow is 70 ± 14 ka (Kirby et al., 2008) and suggest late Pleistocene throw rates of ~0.2 mm/yr. Thus, the Red Mountain fault appears to exhibit similar slip rates to the Fish Springs fault. From this vantage, one can see the Birch Mountain fault, the northernmost segment of the Sierra Nevada frontal fault system. The scarp is apparent south of Tinemaha Creek, where it displaces steep talus and debris-cones at the base of the range, and north of the creek, as it trends upslope toward the prominent cliff at the base of Birch Mountain. The fault displaces Tiogaage moraines in both the Tinemaha and Birch Creek drainages (Fig. 17). Total throw on the structure in these localities varies from 7 to 9 m, and 36Cl ages of boulders from the moraine crest in Tinemaha Creek indicate an age of 13–15 ka (Greene et al., 2007). Thus, throw rates on this segment of the Sierra Nevada frontal fault are ~0.5–0.7 mm/yr Notably, these rates are two to three times greater than those measured along the southern segments of this fault system (Le et al., 2007). Day 3: Big Pine, Furnace Creek, Indian Creek

Stop 10: Owens Valley Fault at Big Pine Dump Directions. This stop is located at UTM 385102E, 4111923N. Just before entering Big Pine from the south on US-395, turn west (left) on Big Pine Dump Road. At the trash/recycling facility, stay to the left onto the dirt road. Stay straight (west) until the intersection with a power-line road. From here, you will see the old town dump. Drive just past the old dump and park under the power lines (~0.3 mi from start of dirt road). Watch out for old nails. Description. At this stop, we will examine evidence for long-term lateral displacement along the Owens Valley fault. From the parking lot near the recycling center, walk west along

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the dirt road for ~300 m. The high-tension power lines visible to the west mark the approximate trace of the Owens Valley fault through this locality. The E-facing scarp of the fault within basalt ﬂows from Crater Mountain is visible as a ~5–7 m high, broken cliff (Fig. 28). At the base of this scarp, ﬁne-grained alluvium has ﬁlled a small extensional step in the fault zone. Anthropogenic activity (the historic town dump) has obliterated any trace of the 1872 rupture through this alluvium, although a strand of the fault is visible as a small scarp ~300 m north of this locality (Beanland and Clark, 1994). The apparent right-lateral separation of the ﬂow margin with the modern alluvium is interpreted to reﬂect dextral displacement across the fault (Kirby et al., 2008). Subsurface surveys using ground-penetrating radar do not reveal any shallow reﬂectors associated with a buried ﬂow margin on the east side of the fault. Rather, the ﬂow margin visible on the surface appears to represent the former extent of the ﬂow (Kirby et al., 2008). West of the fault, however, young alluvial material did bury the ﬂow margin; prominent islands of basalt are visible above alluvium (Fig. 16). Surveys in this region suggest that the ﬂow margin extends to a position near the prominent protrusion of basalt along the fault scarp (Fig. 16). Restoration of this margin with the exposed ﬂow on the east side of the fault suggests ~235 ± 15 m of lateral displacement along the Owens Valley fault in the past 70 ± 14 k.y. Thus, right-lateral slip rates along the northern Owens Valley fault appear to have been 3.6 ± 1.0 mm/yr over the past 56–80 k.y. (Kirby et al., 2008). Stop 11: Fish Lake Valley Fault at Furnace Creek Directions. This stop is located at UTM 410935E, 4158364N. From Big Pine, take CA-168 east over Westgaard Pass. Continue on CA-168 through Deep Springs Valley to Fish Lake Valley. At the intersection of CA-168 and CA-266 in Oasis, turn north on

Figure 28. Photograph taken looking to the west at fault scarps of the Owens Valley fault zone cutting basalt on the northeast side of Crater Mountain near the Big Pine town dump (Stop 10). Peaks along the eastern Sierra Nevada are visible beyond the scarps.

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Figure 29. View to the northwest along the Death Valley–Fish Lake Valley fault zone at Furnace Creek in central Fish Lake Valley (Stop 11). Note the two prominent fault scarps cutting the Furnace Creek alluvial fan. White Mountain Peak is the prominent summit along the skyline ridge.

CA-266 and go ~7.1 mi to a dirt track leading off to the west (just north of White Wolf Canyon Road to the east). Go ~1.5 mi toward the large shutter-ridge. Continue on the dirt track, bearing to the northwest and paralleling the fault, for another 0.3 mi and park near the incised channel (Fig. 18). Walk ~0.1 mi west up the westernmost fault scarp to the apex of the fan. Description. Just south of the canyon mouth at Furnace Creek, the Fish Lake Valley fault is exposed as two parallel NWstriking strands displacing a late Pleistocene alluvial fan complex (Fig. 29; Frankel et al., 2007b; Reheis et al., 1995; Reheis and Sawyer, 1997). Prominent scarps are exposed in the alluvial fans with the western strand dipping to the NE and the eastern strand dipping to the SW, and with other small normal fault scarps scattered across the fan surfaces (Reheis et al., 1995). The area between the two fault strands is down-dropped in a small pullapart basin (parking area is at the northern end of this pull-apart structure). Just south of the offset late Pleistocene fan is a large NW-SE–striking shutter-ridge, which is bounded on its northeast and southwest sides by the two fault strands (Reheis et al., 1995). Although a number of alluvial fan surfaces are mapped in this location (Reheis et al., 1995), the primary surface of interest is the late Pleistocene Qﬁo deposit (Qﬁ of Reheis et al., 1995), which is displaced by the fault (Fig. 18). Cosmogenic nuclide 10 Be dates from the Qﬁo surface yield an age of 94 ± 11 ka (Frankel et al., 2007b). A recent reexamination of offset channels at this location using ALSM data revised the late Pleistocene offset to 290 ± 20 m (Fig. 18; Frankel et al., 2007b). Combining the 290 ± 20 m of displacement determined from ALSM topography with the cosmogenic 10Be age of 94 ± 11 ka for the Qﬁo surface yields a minimum late Pleistocene, right-lateral slip rate of 3.1 ± 0.4 mm/yr for the Fish Lake Valley fault zone at Furnace Creek (Frankel et al., 2007b).

Stop 12: Fish Lake Valley Fault at Indian Creek Directions. This stop is located at UTM 396137E, 4182771N. From Dyer, Nevada, head ~8.5 mi north on NV-264 (~16.8 mi north of the California-Nevada border). Turn west on Indian Creek Road (dirt) and continue for ~5 mi until you reach the prominent E-facing scarp near the canyon mouth at UTM 395992E, 4183225N. Park just west of the scarp on the south side of the road (Fig. 19). Walk ~0.3 mi south along the base of the scarp to the third deeply incised channel cutting the footwall. Climb up to the top of the scarp at this location for a good view of the fault zone. Description. The fault scarps at Indian Creek are located at the northern end of the northern Death Valley–Fish Lake Valley fault system. The fault zone splays into numerous normal faults in this location (Reheis et al., 1993), however, a signiﬁcant strikeslip component is still present (Fig. 19; Frankel et al., 2007b; Reheis et al., 1993; Reheis and Sawyer, 1997). Most of the faulting at this site displaces the late Pleistocene Qﬁy surface (Figs. 19 and 30; Qﬁ of Reheis et al., 1993). The dextral component of slip at this location is restricted to a single strand of the fault near the eastern range-front of the White Mountains (Fig. 19). The strikeslip component of the fault zone is expressed as a prominent scarp cutting the Qﬁy and Qﬂ surfaces (Frankel et al., 2007b; Reheis et al. 1993; Reheis and Sawyer, 1997). The Qﬁy surface at Indian Creek has a similar set of soil and morphologic characteristics as the Qﬁo surface at Furnace Creek. Eight tightly clustered cosmogenic nuclide 10Be surface exposure dates from boulders on the Qﬁy fan surface have a mean age and standard deviation of 71 ± 8 ka (Frankel et al., 2007b). Frankel et al. (2007b) used ALSM data to revise the late Pleistocene displacement history to 178 ± 20 m on the basis of six offset channels (Fig. 19). A late Pleistocene slip rate of 2.5 ± 0.4 mm/yr results from the offset determined with ALSM

Active tectonics of the eastern California shear zone

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Figure 30. Photograph looking southwest along the right-lateral oblique fault scarp at Indian Creek in Fish Lake Valley at the northern end of the Death Valley–Fish Lake Valley fault zone (Stop 12). The eastern White Mountains range front is visible beyond the scarp. Numerous normal faults displace the distal portion of the Indian Creek; all dextral deformation at this location is accommodated on the pictured fault.

data combined with the 71 ± 8 ka 10Be age of the Qﬁy surface at Indian Creek (Frankel et al., 2007b). ACKNOWLEDGMENTS Research presented in this guidebook was supported by the National Science Foundation (NSF), the Geothermal Program Ofﬁce of the China Lake Naval Air Warfare Center, the Southern California Earthquake Center (SCEC), the National Aeronautics and Space Administration, Lawrence Livermore National Laboratory, The Geological Society of America, and the University of California White Mountain Research Station. SCEC is funded by NSF Cooperative Agreement EAR-0106924 and U.S. Geological Survey Cooperative Agreement 02HQAG0008. This is SCEC contribution 1131. ALSM data were collected by the National Center for Airborne Laser Mapping at the University of Florida. A number of people contributed to the work presented in this ﬁeld guide, including S. Briggs, D. Burbank, N. Dawers, J. Helms, E. Hauksson, J. Hoeft, B. Miller, M. Reheis, M. Rogers, G. Roquemore, T. Sheehan, D. Slemmons, and R. Whitmarsh. Thoughtful reviews by Ernie Duebendorfer and Nathan Niemi helped improve the clarity and presentation of this ﬁeld trip guidebook. REFERENCES CITED Argus, D.F., and Gordon, R.G., 1991, Current Sierra Nevada–North America motion from very long baseline interferometry: implications for the kinematics of the western United States: Geology, v. 19, p. 1085–1088, doi: 10.1130/0091-7613(1991)019<1085:CSNNAM>2.3.CO;2. Argus, D.F., and Gordon, R.G., 2001, Present tectonic motion across the Coast Ranges and San Andreas fault system in central California: Geological Society of America Bulletin, v. 113, p. 1580–1592, doi: 10.1130/00167606(2001)113<1580:PTMATC>2.0.CO;2. Atwater, T., 1989, Plate tectonic history of the northeast Paciﬁc and western North America, in Winterer, E.L., Hussong, D.M., and Decker, R.W., eds.,

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Stevens, C.H., and Greene, D.C., 1999, Stratigraphy, depositional history, and tectonic evolution of Paleozoic continental-margin rocks in roof pendants of the eastern Sierra Nevada, California: Geological Society of America Bulletin, v. 111, p. 919–933, doi: 10.1130/0016-7606(1999)111<0919:SDHATE> 2.3.CO;2. Stevens, C.H., Stone, P., Dunne, G.C., Greene, D.C., Walker, J.D., and Swanson, B.J., 1997, Paleozoic and Mesozoic evolution of east-central California: International Geology Review, v. 39, p. 788-829. Stevens, C.H., Stone, P., Dunne, G.C., Greene, D.C., Walker, J.D., and Swanson, B.J., 1998, Paleozoic and Mesozoic evolution of east-central California, in Ernst, W.G., and Nelson, C.A., eds., Integrated Earth and environmental evolution of the southwestern United States, The Clarence A. Hall, Jr., Volume: Boulder, Colorado, Geological Society of America International Book Series 1, p. 119–160. Stevens, C.H., Stone, P., and Greene, D.C., 2003, Correlation of Permian and Triassic deformations in the western Great Basin and eastern Sierra Nevada: Evidence from the northern Inyo Mountains near Tinemaha Reservoir, east-central California: Geological Society of America Bulletin, v. 115, p. 1309–1311, doi: 10.1130/B25385D.1. Stewart, J.H., 1967, Possible large right-lateral displacement along fault and shear zones in the Death Valley–Las Vegas area, California and Nevada: Geological Society of America Bulletin, v. 78, p. 131–142, doi: 10.1130/0016-7606(1967)78[131:PLRDAF]2.0.CO;2. Strane, M.D., 2007, Slip rate and structure of the nascent Lenwood fault zone, Eastern California [M.S. Thesis]: Chapel Hill, University of North Carolina, 55 p. Stockli, D.F., Dumitru, T.A., McWilliams, M.O., and Farley, K.A., 2003, Cenozoic tectonic evolution of the White Mountains, California and Nevada: Geological Society of America Bulletin, v. 115, p. 788–816, doi: 10.1130/0016-7606(2003)115<0788:CTEOTW>2.0.CO;2. Sullivan, W.A., and Law, R.D., 2007, Deformation path partitioning within the transpressional White Mountain shear zone, California and Nevada: Journal of Structural Geology, v. 29, p. 583–599, doi: 10.1016/j.jsg.2006.11.001. Thatcher, W., Foulger, G.R., Julian, B.R., Svarc, J., Quilty, E., and Bawden, G.W., 1999, Present-day deformation across the Basin and Range province, western United States: Science, v. 283, p. 1714–1718, doi: 10.1126/ science.283.5408.1714. Thompson, R.A., Milling, M.E., Fleck, R.J., Wright, L.A., and Roger, N.W., 1993, Temporal, spatial, and compositional constraints on volcanism associated with large-scale crustal extension in central Death Valley, California: Eos (Transactions, American Geophysical Union), v. 74, p. 624. Troxel, B.W., 1994, Right-lateral offset of ca. 28 km along a strand of the southern Death Valley fault zone, California: Geological Society of America Abstracts with Programs, v. 26, no. 6, p. 99. Unruh, J.R., Hauksson, E., Monastero, F.C., Twiss, R.J., and Lewis, J.C., 2002, Seismotectonics of the Coso Range–Indian Wells Valley region, California: Transtensional deformation along the southeastern margin of the Sierran microplate, in Glazner, A.F., Walker J.D., and Bartley, J.M., eds., Geologic evolution of the Mojave Desert and Southwestern Basin and Range: Geological Society of America Memoir 195, p. 277–294. Unruh, J.R., Humphrey, J., and Barron, A., 2003, Transtensional model for the Sierra Nevada frontal fault system, eastern California: Geology, v. 31, p. 327– 330, doi: 10.1130/0091-7613(2003)031<0327:TMFTSN>2.0.CO;2. Unruh, J.R., Monastero, F.C., and Pullammanappallil, S.K., 2008, The nascent Coso metamorphic core complex, east-central California: Brittle upper plate structure revealed by reﬂection seismic data: International Geology Review (in press). Vines, J.A., 1999, Emplacement of the Santa Rita Flat pluton and kinematic analysis of cross-cutting shear zones, eastern California [M.S. Thesis]: Blacksburg, Virginia Polytechnic Institute and State University, 89 p. Walker, J.D., and Whitmarsh, R.W., 1998, A tectonic model for the Coso geothermal area: U.S. Department of Energy Proceedings Geothermal Program Review XVI, April 1–2, Berkeley, California, p. 2-17–2-24. Weldon, R., Scharer, K., Fumal, T., and Biasi, G., 2004, Wrightwood and the earthquake cycle: What a long recurrence record tells us about how faults work: GSA Today, v. 14, no. 9, p. 4–10, doi: 10.1130/10525173(2004)014<4:WATECW>2.0.CO;2. Wells, S.G., McFadden, L.D., and Dohrenwend, J.C., 1987, Inﬂuence of late Quaternary climatic changes on a desert piedmont, eastern Mojave desert, California: Quaternary Research, v. 27, p. 130–146, doi: 10.1016/00335894(87)90072-X. Wernicke, B., Axen, G.J., and Snow, J.K., 1988, Basin and Range extensional tectonics at the latitude of Las Vegas, Nevada: Geological Society of America

Active tectonics of the eastern California shear zone Bulletin, v. 100, p. 1738–1757, doi: 10.1130/0016-7606(1988)100<1738: BARETA>2.3.CO;2. Wernicke, B., Davis, J.L., Bennett, R.A., Normandeau, J.E., Friedrich, A.M., and Niemi, N.A., 2004, Tectonic implications of a dense continuous GPS velocity ﬁeld at Yucca Mountain, Nevada: Tectonics, v. 109, doi: 10.1029/2003JB002832. 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. Wesnousky, S.G., and Jones, C.H., 1994, Oblique slip, slip partitioning, spatial and temporal changes in the regional stress ﬁeld, and the relative strength of active faults in the Basin and Range, western United States: Geology, v. 22, p. 1031–1034, doi: 10.1130/0091-7613(1994)022<1031: OSSPSA>2.3.CO;2. Whistler, D.P., and Burbank, D.W., 1992, Miocene biostratigraphy and biochronology of the Dove Spring Formation, Mojave Desert, California, and characterization of the Clarendonian mammal age (late Miocene) in California: Geological Society of America Bulletin, v. 104, p. 644–658, doi: 10.1130/0016-7606(1992)104<0644:MBABOT>2.3.CO;2.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 18 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 11 2008

Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic–Paleozoic boundary Stephen M. Rowland Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154-4010, USA Lynn K. Oliver Inyo National Forest, 798 North Main Street, Bishop, California 93615, USA Melissa Hicks ExxonMobil Upstream Research Company, P.O. Box 2189, GW3 966B Houston, Texas 77252-2189, USA

ABSTRACT Esmeralda County, Nevada, is extraordinary for the presence of Ediacaran and early Cambrian reefs at several stratigraphic positions. In this road log and ﬁeld guide we present descriptions and interpretations of the most instructive exposures of three of these reef-rich intervals: (1) the Mount Dunfee section of the Middle Member of the Deep Spring Formation (Ediacaran in age), (2) the Stewart’s Mill exposure of the Lower Member of the Poleta Formation (mid-early Cambrian), and (3) an exposure on the north ﬂank of Slate Ridge of reefs near the top of the Harkless Formation (latest early Cambrian). We introduce the term “congruent ecosystems” for ecosystems of different age that occupied similar environments. The Ediacaran reefs of the Deep Spring Formation and the early Cambrian reefs of the Lower Member of the Poleta Formation occupied similar environments but exhibit distinctively different ecological structure. Thus we propose these two reef complexes as our premier example of non-congruent communities within congruent ecosystems. Keywords: Cambrian reefs, Ediacaran reefs, congruent ecosystems, archaeocyaths, stromatolites INTRODUCTION

was such a favorable locality for reef development at that time doubtless involves a combination of paleogeographic factors, including latitude, ocean currents, wind currents, and the conﬁguration of the shelf margin (Fig. 1). The following four stratigraphic units all contain reef-rich intervals: (1) the Middle Member of the Deep Spring Formation, (2) the Montenegro Member of the Campito Formation,

Esmeralda County, Nevada, USA, is an extraordinary region for examining reefs of Ediacaran and early Cambrian age. Probably nowhere else in the world has a better representation of accessible, well-preserved, well-exposed reefs that formed at multiple stratigraphic positions within this time interval. The reason this

Rowland, S.M., Oliver, L.K., and Hicks, M., 2008, Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic–Paleozoic boundary, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 83–100, doi: 10.1130/2008.ﬂd011(04). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Ediacaran and Early Cambrian reefs of Esmeralda County

Figure 1. Paleogeographic setting of the Ediacaran and early Cambrian reefs of Esmeralda County. Modiﬁed from Hicks (2006a).

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(3) the Lower Member of the Poleta Formation, and (4) the Harkless Formation (Fig. 2). Of these four “reefy” intervals, the best-known and most thoroughly studied examples of three of them occur quite close to one another in Esmeralda County. The exception is the Montenegro Member of the Campito Formation; the best-known Montenegro reef occurs in the White Mountains of Inyo County, California (Fuller, 1976; Morgan, 1976; Zhou, 1995), which makes it impractical for us to examine on this short ﬁeld trip. During the past half century, the Ediacaran and Cambrian strata of Esmeralda County, Nevada, and adjacent Inyo County, California, have played an important role in the sorting out of biological and sedimentological events across the ProterozoicPhanerozoic boundary, within North America and globally. This history was recently summarized by Rowland and Corsetti (2002). A ﬁeld trip to two of the stops visited on this trip (Stops 1 and 2.2), as well as to stops not included in this ﬁeld guide, was conducted in 2005 by Anderson et al. (2005), with an emphasis on microbialites. Their guidebook is recommended for researchers interested in the reefs described in this ﬁeld guide, and in microbialites of the Neoproterozoic and Cambrian in general. OBJECTIVES OF THIS FIELD TRIP The objectives of this ﬁeld trip are (1) to provide an opportunity for participants to examine reefs of the Deep Spring, Poleta, and Harkless Formations; (2) to stimulate discussion among participants—at the outcrop and around the campﬁre— about reasons for similarities and differences in these reefs; and (3) to explore with participants the concept of “congruent ecosystems” on opposite sides of the Proterozoic-Phanerozoic boundary. We also invite discussion about the question of why the early Cambrian experiment in reef-building by metazoans

was so short lived. The consortium of reef-building archaeocyaths and calcimicrobes existed on Earth for just 11 m.y., from the base of the Tommotian Stage to the top of the Toyonian Stage (Fig. 3). This was followed by an interval of ~40 m.y. (middle Cambrian through the Early Ordovician), during which virtually no metazoan-built reefs developed anywhere on Earth. Several hypotheses have been proposed for this metazoan-reef– free interval (summarized in Rowland and Shapiro, 2002). To us, the most attractive hypotheses are those that involve phenomena associated with global climate change, but these are difﬁcult to rigorously test. We are hoping for a lively discussion of this topic on this trip. WHAT DO WE MEAN BY “CONGRUENT ECOSYSTEMS”? In this ﬁeld guide we use the term “congruent ecosystems” to characterize ecosystems that occupy the same suite of environments at different times. This is an extension of the term “congruent communities,” originally used by Walker and Laporte (1970). Below is a brief discussion of the history of this terminology and our use of it. Walker and Laporte (1970) coined the term “congruent fossil communities” for fossil communities of different age that occupied similar environments and contain taxa which, although not necessarily closely related taxonomically, have similar autecological characteristics. Their speciﬁc examples were four communities that occupied supratidal, high intertidal, low intertidal, and subtidal carbonate environments in the Ordovician Black River Group of New York and four similar communities that occupied the same environments in the Devonian Manlius Formation, also of New York. Sheehan (1996) adapted the concept of congruent communities to Boucot’s (1983) concept of Ecologic

Ediacaran and Early Cambrian reefs of Esmeralda County

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Wyman 0m Figure 2. Stratigraphic column of Ediacaran and Lower Cambrian of the White-Inyo Range and Esmeralda County region. Modiﬁed from Hicks (2001).

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Figure 3. Spindle diagrams showing diversity of reef-building taxa in the Ediacaran and Cambrian. From Rowland and Hicks (2004).

Evolutionary Units (EEUs). He also renamed Boucot’s EEUs and deﬁned the fauna within each EEU as an “Evolutionary Fauna,” which he characterized as having had “community stasis” during the EEU in which it occurred. As exempliﬁed by the Ordovician and Devonian examples of Walker and Laporte (1970), a community within any of Sheehan’s Evolutionary Faunas is inferred to have had a similar ecologic structure to a community in a similar environmental setting in a different EEU. In Sheehan’s (1996) terminology, the Cambrian period has two EEUs, which he called C1 and C2. C1 corresponds to the portion of the early Cambrian in which archaeocyaths occur (Tommotian through Toyonian Siberian Platform Stages; Fig. 3). On this ﬁeld trip, this EEU includes the Poleta and Harkless Formations. Sheehan (1996) suggested that it may be necessary to erect another EEU in the pre-trilobite portion of the Cambrian, the Nemakit-Daldynian interval of the Siberian Platform (Fig. 3). He also proposed an EEU for the late Neoproterozoic, when the Ediacaran fauna lived. He named this unit E1. On this ﬁeld trip, we will examine reefs that occur in Sheehan’s EEUs E1 and C1. However, the structure of the reefbuilding communities in these two intervals is very different. In contrast to the similarities observed between EEUs within the Phanerozoic eon, such as the Ordovician and Devonian examples of Walker and Laporte (1970), a major point of this ﬁeld trip will be that reef communities across the EdiacaranCambrian boundary exhibit very different ecological structure, although they lived in comparable environments. For this reason, we refer to these as “non-congruent communities” within “congruent ecosystems.”

STOP 1: THE STROMATOLITE REEF COMPLEX OF THE MIDDLE DEEP SPRING AT MOUNT DUNFEE (EDIACARAN IN AGE) At Mount Dunfee, near Gold Point, Nevada, the Middle Member of the Deep Spring Formation is unusually rich in stromatolites and other microbialites. The Neoproterozoic-Cambrian boundary at this locality, as deﬁned by the lowest occurrence of Treptichnus pedum, occurs high within the Middle Member, above the highest of the microbialite horizons. This section thus provides an excellent opportunity to examine morphological diversity in latest Neoproterozoic reefs. Figure 4 is a stratigraphic column of this section, while Figure 5 shows two photographs of the section. We recognize three depositional systems within the Middle Member Deep Spring Formation at Mount Dunfee. The ﬁrst is a siliciclastic intertidal and shallow subtidal system that is represented by quartz siltstone and sandstone in the lower 42 m of the 175m-thick section. These sediments were deposited in tide- and storm-dominated subtidal and intertidal environments prior to the initiation of carbonate sediment on the miogeoclinal ramp. Analysis of herringbone cross-stratiﬁcation within the dolo-allochemic quartz sandstone shows a bimodal pattern that records tidal ﬂow perpendicular to the paleo-shelf margin. Overlying these siliciclastic-dominated sediments are oolites and microbialites, representing peritidal reefs and shoals. Strata of this carbonate-dominated depositional system occupy two intervals within the section. They cap the siliciclastic-dominated interval at the bottom of the section, and they also occupy most of the

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Figure 4. Stratigraphic column of the Middle Member of the Deep Spring Formation at Mount Dunfee. From Oliver (1990).

upper half of the section. In each case, these reef-shoal sediments complete a shoaling-upward cycle within the section. The third depositional system is a mixed siliciclastic-carbonate shallow subtidal system, represented by a 40-m-thick interval of micritic sandstone roughly in the middle of the section. These sediments record the early phase of a second shoaling-upward cycle within the Middle Member, during a time of greater carbonate production. Figure 6 is a stratigraphic column of a portion of the Middle Member, showing microbial horizons A through E and adjacent lithologies. The microbialites are morphologically and structurally diverse, including bioherms of digitate stromatolites, bioherms and biostromes of inclined columnar stromatolites, isolated massive and hemispheroidal stromatolites, stromatolitic thrombolites, and a biostrome of cryptomicrobial boundstone. Some of these were microbial reefs that had topographic relief, formed in active agitated waters, and exerted a physical control over their environment. A few stromatolitic thrombolites occur in this section. This locality is the oldest known occurrence of thrombolites in southwestern North America, although considerably older

examples are known from northwestern Canada. Figure 7 shows examples of stromatolites, thrombolites, and oolites. Meter-scale cyclicity is conspicuous within the portion of the section that contains microbial horizons C, D, and E (Fig. 8). The Mount Dunfee microbialites incorporate up to 32 wt% detrital quartz within their microstructures. They represent an intermediate form between “pure” siliciclastic and “pure” carbonate stromatolites—theoretical end members in a continuum. An important prerequisite for preserving microbialite fabrics in the stratigraphic record is early carbonate cementation. The well-preserved, quartz-rich Mount Dunfee microbialites provide a datum concerning the amount of siliciclastic material a microbialite can contain without a substantial loss of preservation potential and morphological variability. Possible close modern analogs of some of the Mount Dunfee stromatolites occur in subtidal, current-swept channels in the Bahamas. Like their ancient counterparts in the Deep Spring Formation, the Bahamian stromatolites are predominantly columnar, they are inclined into the ﬂood tide, and they are closely associated with oolite sand waves that episodically bury them.

Figure 5. (A) View toward the east of the Middle Member Deep Spring section at Mount Dunfee; ~150 m of strata are in view, and beds dip ~40°NE. The Ediacaran-Cambrian boundary is located a few tens of meters below the top of the Middle Member, above Microbial Horizon I, based on the lowest occurrence of the trace fossil Treptichnus pedum at that horizon. (B) View toward the southeast of a portion of the Middle Deep Spring containing microbial horizons A through E. Compare with Figure 6. Strata within the white box are ~12 m thick. Modiﬁed from Oliver and Rowland (2002).

Figure 6. Detailed stratigraphic column of the interval shown in Figure 5B. From Oliver (1990).

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Figure 7. Photographs of exposures in the Mount Dunfee section. (A) Vertical cliff-face exposure of inclined columnar stromatolites; lower microbial interval. (B) Vertical cliff-face exposure of inclined columnar stromatolites; lower microbial interval; irregular shapes of stromatolites is caused by close packing; width of photo is ~35 cm. (C) Inclined columnar stromatolite with parallel margins; lower microbial interval; pocket knife in center of photo is 9 cm long. (D) Lynn Oliver sitting on Bioherm C.2, a lenticular bioherm of inclined stromatolites. (E) Bioherm C.2, a lenticular bioherm of inclined stromatolites; hammer for scale. (F) Columnar stromatolites below, abruptly overlain by oolite; Microbial Horizon D. (G) Stromatolitic thrombolite, which has a mixture of laminated and clotted fabrics; Microbial Horizon H. From Oliver and Rowland (2002).

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Figure 8. Field sketch of ﬁve meter-scale cycles within microbial horizons C, D, and E. Each cycle begins with a stratiform stromatolite, which usually grades abruptly into columnar stromatolites. In cycle 3, oolite grainstone interrupts the normal succession. We interpret these cycles to be deepening upward cycles.

As represented schematically in Figure 9, we suggest that the Mount Dunfee section, with its conspicuous and episodically persistent peritidal reef and shoal facies, represents the nucleation of carbonate sedimentation in each of two shoaling-upward, carbonate-capped cycles within the Middle Member Deep Spring Formation. From this position on the ramp, carbonate sedimentation expanded seaward and cratonward, ultimately building into a narrow, discontinuous, episodically emergent reef and shoal complex. The Mount Dunfee section thus represents a geographically restricted, temporally persistent locus of carbonate sedimentation and microbialite reef construction during the waning portion of the Ediacaran Period. STOP 2.1: THE EARLY CAMBRIAN ARCHAEOCYATHAN— RENALCIS—THROMBOLITE REEF COMPLEX OF THE LOWER POLETA AT STEWART’S MILL The Stewart’s Mill locality of the Lower Member of the Poleta Formation is the most instructive exposure of the reef facies of this stratigraphic unit. Aspects of this exposure have been described by Rowland (1984), Rowland and Gangloff (1988), and Rowland and Shapiro (2002). The exposure is a pyramidshaped hill (Fig. 10). The lower third is green shale of the Mon-

tenegro Member of the Campito Formation. The middle portion of the hill consists of 70 m of thrombolitic and archaeocyathanRenalcis boundstone of the Lower Member of the Poleta Formation (Fig. 11). Capping the hill is a cliff-forming interval, 56 m thick, of oolite, biostromes, and packstone (Figs. 11 and 12). The interval of main interest for this ﬁeld trip is the 70 m “reefy” interval below the cliffs. There is a complex mosaic of interﬁngering facies in this portion of the exposure (Fig. 11), but the basic succession consists of the following facies, in ascending order: (1) bioclastic lime mudstone, (2) thrombolites with sparse archaeocyaths, (3) Renalcis-dominated boundstone, (4) green mudshale with lenses of skeletal wackestone, (5) archaeocyathdominated boundstone, and (6) oolite grainstone. We interpret Facies 4, the green mudshale (which is poorly exposed), to represent a bypass channel by which siliciclastic sediment was shunted across the carbonate shelf margin. Of particular interest within this exposure is the conspicuous ecological zonation that occurs within the succession of reef-lagoon facies. The lower portion of this interval is dominated by thrombolite, with archaeocyaths comprising less than 3% of the volume of the rock (Fig. 13). In stark contrast, the uppermost portion of this interval is dominated by conspicuous branching archaeocyaths, which comprise up to 38% of the volume of the rock (Fig. 14). Figure 15 summarizes the characteristics within each zone. The three most conspicuous changes that occur, from bottom to top within this 65-m-thick interval, are (1) the relative abundance of archaeocyaths increases (Fig. 16); (2) the morphology of archaeocyaths changes, with non-branching forms dominant throughout most of the section and branching forms becoming dominant within the uppermost 30 m (Fig. 14); and (3) decimeter- and meter-scale cavities (preserved as dolomite-rich, orange patches) are absent in the lower one-third of this interval, but are very conspicuous in the upper two-thirds. Such zonation has been described in fossil reefs of a variety of ages, including Ordovician, Silurian, Devonian, and Cretaceous (Walker and Alberstadt, 1975). This locality is the only example of ecological zonation described in a Cambrian reef. Walker and Alberstadt (1975) interpreted such zonation to be a record of ecological succession, but in cases such as this one, where very long time intervals are involved, ecological succession is not the process being recorded in the rocks (see discussion by Rowland and Gangloff, 1988). Rather, we interpret the conspicuous zonation in the Lower Poleta at Stewart’s Mill to represent adjacent facies that migrated over one another during a marine transgression. Using these three trends, Rowland and Shapiro (2002) divided this interval into three environmental zones: (1) a lower back-reef lagoon zone, (2) a middle low-energy reef-crest zone, and (3) a high-energy reef-crest zone (Fig. 15). The presence of conspicuous primary cavities (now ﬁlled with partially dolomitized micrite) distinguishes the reef-crest facies from the underlying lagoonal facies, and the presence of abundant, conspicuous branching archaeocyaths distinguishes the high-energy reef-crest zone from the underlying low-energy reef-crest zone.

Ediacaran and Early Cambrian reefs of Esmeralda County

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l tida Sub ow l l a h l ed S ida Mix er t Int d e at dal min bti -do w Su d n o Sa hall and and S Silt

elf Sh e Fac lf ore She l/Sh a pen d O i tert ated w In min al d Lo -do e d t n a rtid in Sa Inte om d d d te San ina om d-d u M

Portion of Lower Member Wood Canyon Formation

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Base of Middle Member Deep Spring Formation Intertidal & Supratidal Carbonates Ediacaran - Cambrian Boundary (lowest occurrence of Treptichnus pedum)

Figure 9. Schematic regional synthesis of Middle Member Deep Spring facies at Mount Dunfee and correlative strata to the northwest and southeast.

Figure 10. Photograph of the Stewart’s Mill exposure of the Lower Member of the Poleta Formation and uppermost portion of the Montenegro Member of the Campito Formation. Lower third of hillside is shale of the uppermost Campito Formation. Middle third is 70-m-thick interval of predominantly archaeocyathan-Renalcis reef facies, interrupted by a poorly exposed shale and siltstone interval. Cliff-forming upper third of exposure is 60-m-thick interval, predominantly oolite. View is toward the north. Compare with Figure 11. Photo by S. Rowland.

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Figure 11. Lower Poleta facies at the Stewart’s Mill locality; qtz siltst—quartz siltstone. From Rowland (1981).

It is noteworthy that the Lower Poleta reef complex at Stewart’s Mill has abundant microbialites but, in stark contrast to the Deep Spring reefs at Mount Dunfee, there are no stromatolites. All of the microbialite in the Poleta at this locality has a clotted, thrombolitic texture.

Poleta reefs and oolite shoals formed during an interval of carbon isotopic stasis.

Chemostratigraphy of the Lower Poleta

STOP 2.2: PATCH REEFS NEAR THE TOP OF THE HARKLESS FORMATION ON SLATE RIDGE: THE LAST GASP OF METAZOAN REEF-BUILDING IN THE EARLY CAMBRIAN

As part of Hicks’s dissertation research, we sampled this exposure of the Lower Poleta for isotopic analysis, with the hope of using chemostratigraphy to assist in intercontinental correlation. Her results are presented in Figure 17. Although the early Cambrian is known for its carbon isotope excursions, surprisingly this interval showed very little variation in its δ13C record. There is a slight drift in δ13C values ranging from 0‰ to 1‰ at the bottom of the section to −1‰ at the top. The fact that the δ13C values do not covary with the δ18O values (Fig. 17) suggests that the stasis in the δ13C values is a primary signal and not a product of diagenetic alteration (Hicks, 2006a). Apparently, the Lower

The ﬁnal stop on this ﬁeld trip will be in the upper portion of the Harkless Formation to examine some meter-scale patch reefs studied by Hicks (2001) (Fig. 18). These Harkless reefs are among the youngest Cambrian reefs in the world in which metazoans played a signiﬁcant constructional role. They are approximately equivalent in age to the well-studied reefs of the Forteau Formation of eastern Canada. Following the disappearance of these late early Cambrian archaeocyath-rich reefs, microbialites (stromatolites, thrombolites, and dendrolites) became the prevailing reefs through the middle and late Cambrian (Furongian) and Early Ordovician (Rowland and Shapiro, 2002).

Figure 12. Stratigraphic column of the oolite shoal facies complex at the Stewart’s Mill locality. Zone numbers refer to inferred amount of agitation in the depositional environment; zone 1 is constant agitation, while zone 4, at the other extreme, represents only occasional, storm-generated agitation. From Rowland (1981).

Figure 14. Abundant branched archaeocyaths high in the Lower Poleta boundstone interval at Stewart’s Mill. Archaeocyaths comprise up to 38% of the volume of the rock at this level. Photo by S. Rowland.

Figure 13. Typical exposure of thrombolite-rich interval low in the Lower Poleta section at Stewart’s Mill. Archaeocyaths are sparse within this facies, comprising less than 3% by volume. This facies corresponds to the back-reef lagoon zone of Figure 15. Rock hammer for scale. Photo by S. Rowland.

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Figure 15. Vertical ecological zonation in the reef-lagoon facies complex of the Lower Member of the Poleta Formation at Stewart’s Mill. From Rowland and Hicks (2004).

Figure 16. Facies relationships within the reef-lagoon interval at Stewart’s Mill. Numbers below the black triangles indicate the volumetric abundance of archaeocyaths. From Rowland and Gangloff (1988).

Ediacaran and Early Cambrian reefs of Esmeralda County Stewart's Mill Poleta Formation Section -18

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Figure 17. Values of δ18O and δ13C (‰) for the Lower Poleta at Stewart’s Mill. From Hicks (2006a).

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Reef A Figure 18. Three meter-scale patch reefs near the top of the Harkless Formation at Site 1 (Stop 2.2) of Hicks (2001). View is toward the north. Photo by M. Hicks.

Figure 19. Stratigraphic column of an interval near the top of the Harkless Formation, at Site 1 (Stop 2.2) of Hicks (2001), including a patch reef. From Hicks (2001).

Ediacaran and Early Cambrian reefs of Esmeralda County Figure 19 is a stratigraphic column of the locality we will visit, and Figure 20 shows the sedimentological and paleontological details of the reefs and associated facies. As is the case with the Forteau reefs, these Harkless reefs are associated with a high diversity fauna of dwelling organisms that were important components of community structure but were not involved in reef construction. Figure 21 shows the volumetric percentages of various constituents of the reef framestone. Some reefs within the uppermost Harkless contain one of the earliest reported corals from North America, Harklessia yuenglingensis, which has similarities with tabulate corals (Hicks, 2006b). Unfortunately, the reefs with Harklessia are not easily accessible, so these early Cambrian corals will not be observed on this ﬁeld trip. ROAD LOG DAY 1: EDIACARAN MICROBIAL REEFS OF THE MIDDLE MEMBER DEEP SPRING FORMATION AT MOUNT DUNFEE This day will involve a drive of ~200 mi (320 km) into Esmeralda County, Nevada, USA. We’ll travel north from Las Vegas on U.S. 95, with a stop for lunch and a restroom break in Beatty, Nevada. The road distance from the University of Nevada–Las Vegas (UNLV) to Beatty is ~125 mi (200 km). The detailed road log begins at the Death Valley Nut and Candy Company in Beatty. We’ll spend about three hours in the afternoon examining the microbial reefs of the Horse Spring Formation. Cumulative mi (km) 0.0

(0.0)

51.9

(83.5)

58.9

(94.7)

66.4 (106.8)

66.5 (106.9) 66.7 (107.3) 67.0 (107.8)

68.2 (109.7) 69.2 (111.3) 69.6 (112.0)

Directions Leave the Death Valley Nut and Candy Company parking lot (and Eddy World gas station) and head north (left) on U.S. 95. Turn left (southwest) on Nevada Highway 266 at Lida Junction. Turn left (southwest) on Nevada Highway 744 toward Gold Point. “Downtown” (Gold St. and 2nd Ave.) Gold Point, a semi-ghost town. Turn left on 2nd Avenue. Bear right on main dirt road (extension of Orleans St.). Turn right at intersection of two dirt roads. Turn left on dirt road. (On the Gold Point 7.5 min Quadrangle map, this is a prominent road that heads southeast around the south side of Mount Dunfee.) Junction; no turn. Junction marked by a rusty old street sign; no turn. Bold ridge of rhyolite on left (the “rhyolite mammoth”); park on left side of road.

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Stop 1: Microbialite Reefs and Associated Sedimentological Features of the Middle Member of the Deep Spring Formation Here we begin a hike of ~1.5 mi, over uneven terrain, into the Mount Dunfee section. The destination is the exposure beneath the second “e” in the word “Dunfee” on the quadrangle map. See detailed description above. Cumulative mi (km)

Directions

69.6 (112.0) Leave the “rhyolite mammoth” and return to Gold Point. 72.8 (117.1) Turn right on Main Street (Mitchell’s Mercantile is a prominent building on Main St.). 72.9 (117.3) Turn left on dirt road that leads north across Lida Valley 80.0 (128.7) Turn right at intersection of two dirt roads. 80.4 (129.4) Intersection with Nevada Highway 266; turn right onto pavement. 81.4 (131.0) Turn left onto well-used dirt road marked by a faded, trapezoid-shaped Bureau of Land Management sign just beyond turn. 84.3 (135.6) Bear left at fork in road. 84.4 (135.8) Turn left onto a track that leads into the hills. 85.0 (136.8) Campsite in abandoned quarry. DAY 2: EARLY CAMBRIAN ARCHAEOCYATHANRENALCIS REEFS OF THE POLETA AND HARKLESS FORMATIONS We’ll visit two localities this morning, both of which are just a few miles from our campsite. The ﬁrst will be the massive reef and shoal complex at the Stewart’s Mill locality of the Lower Member of the Poleta Formation, and the second will be some meter-scale patch reefs near the top of the Harkless Formation. Cumulative mi (km)

Directions

85.0 (136.8) Leave campsite and retrace route back to Nevada Highway 266. 88.6 (142.6) Turn right (west) onto Nevada Highway 266. 89.6 (144.2) Turn left onto dirt road (marked by a stop sign). 89.9 (144.6) Cross intersection; no turn. 91.0 (146.4) Turn right onto track that leads toward prominent pyramid-shaped hill. 91.1 (146.6) Parking area. Stop 2.1: Stewart’s Mill Locality of the Lower Member of the Poleta Formation It is a short but steep hike up the hill, through the shales of the Montenegro Member of the Campito Formation, and into

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Figure 20. Sedimentological features of the reef-bearing interval of the Harkless Formation at Site 1 (Stop 2.2). From Hicks (2001).

Figure 21. Mean percentages of constituents in Site 1 (Stop 2.2) reefs. “Other” category refers to brachiopods and unidentiﬁable shell material. Sm. block and lg. block refer to calcite cement. From Hicks (2001).

Ediacaran and Early Cambrian reefs of Esmeralda County the Lower Member of the Poleta Formation. See earlier detailed description of this locality. A Note about Collecting Samples at This Site This exposure of the Lower Poleta is one of the best examples of early Cambrian archaeocyathan reefs in the world. Professors from many universities bring their students here, and paleontologists from all over the world come here to examine this reef complex. If you want to collect samples, please conﬁne your collecting to loose material on the slope, of which there is plenty. Please leave the in situ fossils for future generations of geologists and paleontologists to examine and study in context. Cumulative mi (km) 91.1 91.2 92.3 92.6

(146.6) (146.7) (148.5) (149.0)

101.1 (162.7) 101.7 (163.6)

101.9 (164.0)

103.5 (166.5) 103.8 (167.0)

Directions Retrace route back to well-used dirt road. Turn left on road. Cross intersection; no turn. Junction with Nevada Highway 266. Turn right (east). Turn right (toward Gold Point) onto Nevada Highway 774. As paved road turns to the right, continue straight onto dirt track. (Beginning at this point the route requires high clearance and a low center of gravity; 4-wheel drive recommended.) Cross intersection with another track. No turn. Our route heads toward prominent black peak in the distance. Turn left out of wash onto track that leads up onto desert surface above level of wash. Parking area for Stop 2.2.

Stop 2.2: Patch Reefs of the Harkless Formation Three meter-scale patch reefs of the Harkless Formation are exposed on the low ridge immediately to the north of this parking area. See earlier detailed description of this locality. After examining these reefs, we will eat lunch and then begin our return to Las Vegas. Cumulative mi (km) 103.8 105.9 106.5 113.5 165.4

(167.0) (170.4) (171.4) (182.6) (266.1)

Directions Retrace route back to Nevada Highway 774. Turn right on Nevada Highway 774. Turn right (east) on Nevada Highway 266. Turn right (south) on U.S. 95 toward Beatty. Brief rest stop at the Death Valley Nut and Candy Co. in Beatty. Continue south on U.S. 95 to Las Vegas. Distance from Beatty to UNLV is ~125 mi (200 km). Total distance traveled on this ﬁeld trip is 415 mi (668 km). End of road log.

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ACKNOWLEDGMENTS We thank Tom Anderson, Ernie Duebendorfer, and Russell Shapiro for helpful reviews of this ﬁeld guide, and we thank Becki Huntoon, Peter Starkweather, and Ethan Starkweather for help with ﬁgures. REFERENCES CITED Anderson, T.B., Hicks, M., and Shapiro, R.S., 2005, Microbialite sediments of the Death Valley area, in Stevens, C., and Cooper, J., eds., Western Great Basin Geology: ﬁeldtrip guidebook and volume prepared for the joint meeting of the Cordilleran Section, GSA, and Paciﬁc Section, AAPG: Fullerton, California, Paciﬁc Section SEPM (Society for Sedimentary Geology), p. 67-107. Boucot, A.J., 1983, Does evolution take place in an ecological vacuum? II: Journal of Paleontology, v. 57, p. 1–30. Fuller, D.R., 1976, Paleoenvironmental analysis of an Early Cambrian archaeocyathid reef in the White-Inyo Mountains, California [M.S. thesis]: Idaho State University, 53 p. Hicks, M., 2001, Paleoecology of upper Harkless archaeocyathan reefs in Esmeralda County, Nevada [M.S. thesis]: Las Vegas, University of Nevada, 150 p. Hicks, M., 2006a, Characterizing global archaeocyathan reef decline in the Early Cambrian: Evidence from Nevada and China [Ph.D. dissertation]: Las Vegas, University of Nevada, 138 p. Hicks, M., 2006b, A new genus of Early Cambrian coral in Esmeralda County, southwestern Nevada: Journal of Paleontology, v. 80, p. 609–615, doi: 10.1666/0022-3360(2006)80[609:ANGOEC]2.0.CO;2. Morgan, N., 1976, The Montenegro bioherms: their paleoecology, relation to other archaeocyathid bioherms and to early Cambrian sedimentation in the White and Inyo Mountains, California, in Moore, J.N., and Fritsche, A.E., eds., Depositional Environments of Lower Paleozoic Rocks in the White-Inyo Mountains, Inyo County, California, Paciﬁc Coast Paleogeographic Field Guide 1: Los Angeles, California, Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, p. 13-17. Oliver, L.K., 1990, Stromatolites of the Middle Member of the Deep Spring Formation, Esmeralda County, Nevada [M.S. thesis]: Las Vegas, University of Nevada, 150 p. Oliver, L.K., and Rowland, S.M., 2002, Microbialite reefs at the close of the Proterozoic Eon: The Middle Member Deep Spring Formation at Mount Dunfee, Nevada, in Corsetti, F.A., ed., Proterozoic-Cambrian of the Great Basin and Beyond: Fullerton, California, Paciﬁc Section SEPM (Society for Sedimentary Geology), p. 97–122. Rowland, S.M., 1981, Archaeocyathid bioherms in the Lower Poleta Formation, Esmeralda County, Nevada, in Taylor, M.E., and Palmer, A.R., eds., Second International Symposium on the Cambrian System, Guidebook for Field Trip 1: Denver, Colorado, p. 44-49. Rowland, S.M., 1984, Were there framework reefs in the Cambrian?: Geology, v. 12, p. 181–183, doi: 10.1130/0091-7613(1984)12<181:WTFRIT> 2.0.CO;2. Rowland, S.M., and Corsetti, F.A., 2002, A brief history of research on the Precambrian-Cambrian boundary in the southern Great Basin, in Corsetti, F.A., ed., Proterozoic-Cambrian of the Great Basin and Beyond: Fullerton, California, Paciﬁc Section SEPM (Society for Sedimentary Geology), p. 97–122. Rowland, S.M., and Gangloff, R.A., 1988, Structure and paleoecology of Lower Cambrian reefs: Palaios, v. 3, p. 111–135, doi: 10.2307/3514525. Rowland, S.M., and Hicks, M., 2004, The Early Cambrian experiment in reefbuilding by metazoans, in Lipps, J.H., and Waggoner, B.M., eds., Neoproterozoic-Cambrian biological revolutions: The Paleontological Society Papers, v. 10, p. 107–130. Rowland, S.M., and Shapiro, R.S., 2002, Reef patterns and environmental inﬂuences in the Cambrian and earliest Ordovician, in Kiessling, W., Flűgel, E., and Golonka, J., eds., Phanerozoic Reef Patterns: Tulsa, Oklahoma, SEPM (Society for Sedimentary Geology) Special Publication 72, p. 95–128. Sheehan, P.M., 1996, A new look at Ecological Evolutionary Units (EEUs): Palaeogeography, Palaeoclimatology, Palaeoecology, v. 127, p. 21–32, doi: 10.1016/S0031-0182(96)00086-7.

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Walker, K.R., and Alberstadt, L.P., 1975, Ecological zonation as an aspect of structure in fossil communities: Paleobiology, v. 1, p. 238–257. Walker, K.R., and Laporte, L.F., 1970, Congruent fossil communities from Ordovician and Devonian carbonates of New York: Journal of Paleontology, v. 44, p. 928–944.

Zhou, X., 1995, Lower Cambrian bioherms in central Nevada and eastern California [M.S. thesis]: Las Vegas, University of Nevada, 106 p. MANUSCRIPT ACCEPTED BY THE SOCIETY 24 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 11 2008

Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert Andrew P. Barth Department of Earth Sciences, Indiana University–Purdue University, Indianapolis, Indiana 46202, USA J. Lawford Anderson Department of Earth Sciences, University of Southern California, Los Angeles, California 90089, USA Carl E. Jacobson Department of Earth and Atmospheric Sciences, Iowa State University, Ames, Iowa 50011, USA Scott R. Paterson Department of Earth Sciences, University of Southern California, Los Angeles, California 90089, USA Joseph L. Wooden U.S. Geological Survey, 345 Middleﬁeld Road, Menlo Park, California 94025, USA

ABSTRACT This ﬁeld guide describes a two-and-one-half day transect, from east to west across southern California, from the Colorado River to the San Andreas fault. Recent geochronologic results for rocks along the transect indicate the spatial and temporal relationships between subarc and retroarc shortening and Cordilleran arc magmatism. The transect begins in the Jurassic(?) and Cretaceous Maria retroarc fold-andthrust belt, and continues westward and structurally downward into the Triassic to Cretaceous magmatic arc. At the deepest structural levels exposed in the southwestern part of the transect, the lower crust of the Mesozoic arc has been replaced during underthrusting by the Maastrichtian and/or Paleocene Orocopia schist. Keywords: California, structural geology, petrology, geochronology, tectonics OVERVIEW Achieving the goal of understanding the geodynamic evolution of a convergent continental margin arc requires an understanding of the interplay between magmatic and tectonic processes through time. How did shortening in the southwestern North American Cordillera relate in time and space to arc magmatism?

Advances in geochronology applied to regional ﬁeld and petrologic studies in an exhumed arc and its associated thrust belts in southern California are illuminating the timing of underthrusting and shortening relative to voluminous arc magmatism. Cordilleran foreland shortening and arc magmatism were broadly contemporaneous, but their relative timing remains poorly known. As a result, it has long been argued whether

Barth, A.P., Anderson, J.L., Jacobson, C.E., Paterson, S., and Wooden, J.L., 2008, Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 101–117, doi: 10.1130/2008. ﬂd011(05). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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shortening leads to, or is the structural response to precursory magmatic thickening in the arc. As higher precision U-Pb ages become more readily available for more plutons in the Cordilleran arc, it seems likely that batholith emplacement was an episodic phenomenon superimposed on a longer-term background magmatic ﬂux. In a like manner, more precise geochronologic control on evolution of the foreland fold-and-thrust belt is necessary to better establish the orogen-wide stress ﬁeld over long time scales, and thus clarify whether foreland shortening is a cause or consequence of convergent margin arc magmatism. This ﬁeld trip offers an overview of the tectonic evolution of a portion of the retroarc region and a tilted section through the crust of the Cordilleran arc. An east to west transect will afford us a view of arc tectonics in the Cretaceous, and a top-down view of variations in the composition and emplacement style of Mesozoic igneous rocks in a tilted arc section. Comparison of the plutonic and shortening records allows us to relate deformation in time and space to the progress of arc magmatism. The tilted arc section was likely created during shallow underthrusting of oceanic lithosphere beneath this long-lived Mesozoic arc, and so the timing of this underthrusting event is key to understanding arc extinction and exhumation. In the Orocopia Mountains, exhumation has exposed detached plates of moderate and deep structural levels of the arc. These are underlain by Orocopia Schist, which is part of the Pelona-Orocopia-Rand schist subduction complex underplated beneath southern California and western Arizona during the Laramide orogeny. Here we will consider both the underplating and exhumation history of the schist, as well as geologic relations within the overlying crystalline and supracrustal rocks of native North America. On this trip, we will follow an east to west transect from the retroarc fold-and-thrust belt along the Arizona–California border, westward into the Mesozoic arc. Along this transect, the timing of Mesozoic plutonism is now reasonably well characterized using U-Pb geochronology, which has also been applied to estimating the timing of shortening in the retroarc and subarc regions. At this latitude, the Mesozoic plutonic record incorporates overlapping segments of Permo-Triassic, Middle Jurassic, Late Jurassic and Late Cretaceous arc segments. As we progress further to the west into the core of the arc, we will see a region of the arc that enjoyed shallow underthrusting and consequent extreme extension during and following the early

Cenozoic Laramide orogeny. The result of this variable extension is regional west-side-up tilting that allows us a view of depth-dependent changes in arc magmatism and processes of magma transport and emplacement. DAY 1 Directions to Stop 1 Depart Blythe, traveling west on Interstate 10 (I-10). Exit Mesa Road [UTM E 0710924 N 3721114 (NAD 27 CONUS)]. Continue west on Black Rock Road to Stop 1 [0708035 3721024]. Stop 1. Mule Mountains Thrust Zone and McCoy Mountains Formation This stop ﬁnds us in the south-central part of the McCoy basin, a west-northwest–trending basin in western Arizona and southeastern California. The later structural evolution of the basin is characterized by crystalline thrust sheets of opposing vergence, primarily exposed along the margins of the basin. The basin margins are visible as Middle Jurassic plutonic rocks thrust over Jurassic volcanic rocks in the Mule Mountains visible to the south, and Proterozoic basement and Paleozoic cratonal cover in the Maria Mountains visible to the northeast (Stone, 2006; Fig. 1). At this location, we can see rocks characteristic of the basin and its margins; Jurassic metavolcanic rocks of the Dome Rock sequence are thrust northward over clastic sediments of the Jurassic(?) and Cretaceous McCoy Mountains Formation. The earlier evolution of the McCoy basin is controversial, depending on interpretation of the origin of the 4–7.5 km of clastic sediment of the McCoy Mountains Formation that have yielded very few, problematic fossils. An inferred Jurassic depositional age for the McCoy Mountains Formation led to an early hypothesis that during most of its evolution the basin was bounded by sinistral transtensional faults of the hypothetical Mojave-Sonora megashear (Harding and Coney, 1985; Saleeby and Busby-Spera, 1992; Anderson and Nourse, 2005). Alternatively, an inferred Cretaceous depositional age (Reynolds et al., 1986; Stone et al., 1987; Tosdal, 1990; Tosdal and Stone, 1994) led to the hypothesis that at least some basin sedimentation was associated with shortening in the Maria fold-and-thrust belt.

Figure 1. Simpliﬁed cross section (adapted from Stone, 2006) of the McCoy Mountains Formation (gray shaded) in the McCoy basin. Tu—undifferentiated Tertiary; Mg—Mesozoic granitic and gneissic rocks; Jv—Jurassic volcanic rocks; Pz—Paleozoic and Triassic(?) sedimentary rocks; Xg—Proterozoic crystalline rocks.

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Recent petrography, zircon geochronology and geochemistry of McCoy sandstones provide new limits on the timing and origin of McCoy basin ﬁll (Barth et al., 2004). Detrital zircon populations in sandstones vary systematically with stratigraphic height (Fig. 2) and sandstone composition. Minimum detrital zircon ages of 116 and 109 Ma in the lower part of basal sandstone member 2 demonstrate that >90% of the formation was deposited in post-Aptian, middle Early to middle Late Cretaceous time. Comparison of detrital zircon ages to sandstone compositions indicates that sand was derived from both sides of the basin, and that deposition was synchronous with volcanism to the west and shortening and exhumation in the fold-and-thrust belt. These results support the hypothesis that the McCoy basin originated as a retroarc foreland basin during Cretaceous thrusting. This result is signiﬁcant because the McCoy basin then links shortening and foreland basin sedimentation in the Maria fold-and-thrust belt to contemporaneous retroarc shortening in the southern Sevier belt to the north and the Sonoran thrust belt to the south, recording regional crustal shortening synchronous with voluminous upper crustal arc magmatism. Directions to Stop 2 Depart stop, return east on Black Rock Road to I-10. Continue west on I-10. Exit Eagle Mountain Road [0643527 3730127]. Continue north on Eagle Mountain Road to Stop 2 [0643247 3738425]. Stop 2. Middle Jurassic Eagle Mountain Pluton, Overlain by Pliocene Alkali Basalt

Figure 2. Stratigraphic column for the McCoy Mountains Formation in its type section in the McCoy Mountains (Barth et al., 2004). Symbols to the right of the column show sample locations and U-Pb zircon minimum detrital ages.

Middle to upper crustal portions of two juxtaposed Mesozoic magmatic arcs are exposed in the area being investigated in this ﬁeld trip, one being the Mojave province of the Cordilleran orogen and the other being the “exotic” or “suspect” Tujunga terrane (also termed the “San Gabriel Terrane”), that, in part, comprises the Transverse Ranges, including the San Gabriel Mountains, and which continues southeastward into the Mojave Desert to include desert mountain ranges south of exposures of the McCoy Mountains Formation, including the Chuckwalla, Little Chuckwalla, Chocolate, Pinto, Eagle, and Mule Mountains. The Tujunga terrane, a fault-bounded block underlying 21,000 km2 of southern California and adjacent Arizona, has been termed a suspect terrane due to its distinctive crystalline units that appear to have no correlative ties to native North America, including its exposure of 1190 Ma anorthosite (Barth et al., 2001a). The Tujunga terrane is everywhere allochthonous. Crystallization thermobarometry of dated plutons within the terrane has indicated that much of its apparent “suspect” nature stems from its partial derivation from the middle crust. Batholiths of the central and western Mojave have been studied by Coleman and Walker (1992), Coleman et al. (1992), and Miller and Glazner (1995). Large intrusions also comprise the central and eastern Mojave Desert and early studies include

those of Anderson and Rowley (1981), Beckerman et al. (1982), Miller et al. (1982), Howard et al., (1987), John (1987), Miller et al. (1990), Anderson and Cullers (1990), Miller et al. (1992), Anderson et al. (1992), Young et al., (1992), Miller and Wooden (1994), Gerber et al. (1995) and Mayo et al. (1998). The Mojave Desert region also contains other basement terranes of uncertain origin, including the Joshua Tree terrane. Bender et al. (1993) have concluded that the Joshua Tree terrane is contiguous with the Mojave and that the allochthonous Tujunga terrane represents a displaced, middle crustal section of the Mojave block. Mesozoic plutons of the Tujunga terrane (Barth et al., 1990; Barth et al., 1995) bear strong isotopic and elemental afﬁnities to the plutons of the eastern Mojave, including essentially identical time-transgressive and compositional changes in magmatic arc construction, including both Early Proterozoic and Mesozoic plutons, thus suggesting that the Tujunga terrane is a displaced portion of the Mojave province (Bender et al., 1993; Anderson et al., 1992; Anderson et al., 1993). Mesozoic batholiths of the Mojave Desert and the Tujunga terrane occur in three distinct pulses (Fig. 3). Scattered Triassic intrusions, often distinctly K-feldspar megacrystic, were

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emplaced between 250 and 207 Ma (Barth and Wooden, 2006). After a magmatic lull of ca. 40 Ma, Jurassic magmatism led to widespread intrusion between 165 and 143 Ma. The Independence dike swarm was emplaced at 149 Ma late in this magmatic epoch (Carl and Glazner, 2002). Following another magmatic lull, one of 45 Ma, renewed magmatism in the mid Cretaceous progressed to a peak of activity at 82–72 Ma. Similar-aged magmatic episodes occur throughout the Cordilleran orogen of western North America, but nowhere else in the orogen is there the range of crustal depth of pluton emplacement as seen in the Mojave province, including age-correlative units in the Transverse Ranges. Mid-crustal plutons emplaced at depths greater than 20 km occur both in allochthonous sheets in Mesozoic thrust fault complexes and in the lower plates of Cenozoic metamorphic core complexes (Anderson, 1996). Interestingly, all Mesozoic plutons of the Mojave, regardless of age, were derived from high fO2 magmas, as shown by their high magnetite content. Low fO2 granitic magmas typically have ilmenite as the principle Fe-Ti oxide phase as found for considerable eastern portions of the Peninsular Ranges batholith (Shaw et al., 2003) Triassic Plutonism Several Triassic plutons occur in southern California including in the San Gabriel and Mule Mountains of the Tujunga terrane and the Granite I, Little San Bernardino and San Bernardino Mountains. U-Pb (zircon) age determinations range from 250 to 207 Ma (Barth et al., 1990; Miller, 1977; Frizzell et al., 1986). (Note: there are four named Granite Mountains in the Mojave Desert. What we term Granite I borders the San Bernardino Mountains and Granite II in the central Mojave Desert is adjacent to the Providence Mountains.) The plutons of both the Mojave province and the Tujunga terrane share many common megascopic and petrologic attributes making them much different from those of Jurassic and Cretaceous age. Metaluminous, high K2O + Na2O, and often of low silica (Fig. 4), the plutons are principally of monzonite and quartz monzonite with lesser amounts of monzodiorite and diorite. Most are

alkalic, with a shoshonitic afﬁnity. High abundances of Ba, Sr (>1000 ppm) (Fig. 5), and light rare earth elements are also characteristic as ﬁrst recognized by Miller (1977). Megacrystic K-feldspar is common. High feldspar content is reﬂected by low abundances of Fe, Mg, and Ti. The principal maﬁc minerals are hornblende and biotite ± clinopyroxene and garnet. The garnet coexisting with hornblende in the San Gabriel Mountains Mount Lowe intrusion (Tujunga terrane) and the quartz monzonite pluton of Granite I Mountains (Mojave province), is exceptionally enriched in grossular and andradite components. Barth (1990) has documented that the Mount Lowe was emplaced in the middle crust based on calculated pressures of 5.5–7.0 kbar, consistent with several mineralogical attributes of deep-seated crystallization, including markedly aluminous hornblende (to >11% Al203), calcic garnet, magmatic epidote, and siliceous primary muscovite. For the plutons in the San Bernardino and adjacent Granite I Mountains, Miller (1978) has argued that the magmas were derived from a large ion lithophile element-enriched, quartz eclogite source. The Mount Lowe intrusion is an immense, batholithic-sized (>300 km2), zoned plutonic complex that occurs in the San Gabriel Mountains and in correlative exposures east of San Andreas fault (Chocolate, Little Chuckwalla, Mule, and Trigo Mountains). Based on a broad database of elemental and Sr, Pb, and Nd isotopic data, Barth and Ehlig (1988) and Barth et al. (1990) have argued that the marginal zone of the intrusion was derived from low degrees of partial melting of an eclogitic and enriched subcontinental lithospheric source with virtually no input of continental material, whereas the central zone formed originally in a similar manner but with considerable crustal enrichment. Jurassic Plutonism Magmatic arc construction in the southern Cordillera and in the Tujunga terrane became a major feature of the orogen by the mid-Jurassic. The plutons are largely metaluminous, typically contain hornblende ± clinopyroxene and include gabbro, diorite, quartz monzodiorite, quartz monzonite, quartz syenite, and syenogranite. The granitic rocks are coarse grained and seriate to porphyritic with large, lavender-colored K-feldspar phenocrysts (Tosdal et al., 1989). The plutons have metasomatic effects not encountered in intrusions of Triassic or Cretaceous age. Zones of albitization occur in the Bristol, Ship, and Marble Mountains plutons. Based on stable isotopic data of the altered rocks, Fox and Miller (1990) have interpreted the ﬂuids to be of meteoric origin. Replacement deposits of massive magnetite occur at upper crustal contacts with Paleozoic marbles in the Providence and Eagle Mountains, which have had considerable historical mining interest. Hall et al. (1988) present isotopic data supporting a magmatic origin of the ﬂuids leading to these iron deposits. Crystallization thermobarometry shows a range of determined emplacement depth (Fig. 6). Upper crustal complexes occur in the Providence, Marble, Bristol, Chuckwalla, and Eagle Mountains (1–3 kbar). In contrast, deep-seated Jurassic complexes

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and 206Pb/204Pb deﬁnes a strikingly linear array, suggesting 1.6– 1.7 Ga crustal rocks were an important component of the source region (Anderson et al., 1990; Fox and Miller, 1990; J.L. Wooden, unpublished data). Young et al. (1992) modeled the origin of plutons in the Granite and Bristol Mountains by partial melting of hydrous and enriched mantle coupled with variable assimilativefractional crystallization involving ~10% Proterozoic crust.

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have been identiﬁed the Granite II and Cargo Muchacho Mountains (6–7 kbar) (Anderson, 1996, and references therein). Compositionally, these plutons are unlike those of Cretaceous or Triassic age but no signiﬁcant difference occurs across the inferred position of the Mojave-Tujunga terrane boundary. Total alkalis are elevated, Fe/Mg ratios are intermediate (transitional between calc-alkaline and tholeiitic), and Sr abundances are moderate. Alkali-lime indexes require the existence of both calc-alkalic and alkali-calcic members. Limited isotopic data show δ18O values clustering near 8‰, Sri ranging from 0.706 to 0.709, and εNdt from −6 to −9 (Fig. 7). Covariation of 207Pb/204Pb

Cretaceous Plutonism After extended post-Jurassic magmatic lull, voluminous Cretaceous igneous activity in both the Tujunga terrane and the Mojave region began at ca. 95 Ma and reached a peak at 82– 72 Ma (Miller et al., 1982; Beckerman et al., 1982; Wright et al., 1986, 1987; Anderson et al., 1990). As found for the Jurassic plutons, the depths of emplacement of the Cretaceous vary widely (Anderson, 1996). Mid-crustal intrusions have been found in the San Gabriel, Old Woman, Granite II, Chemehuevi, and Whipple Mountains (Anderson et al., 1988). Plutons in Joshua Tree National Park were largely intruded at shallow crustal levels, but include a mid-crustal section (Needy et al., 2006); upper crustal intrusions have been identiﬁed in the Teutonia and Cadiz Valley batholiths, the Chuckwalla Mountains, and the Sacramento Mountains core complex. Compositionally, and by rock type, the Cretaceous suites of the Mojave and the Tujunga terrane are distinct from the magmatic activity of earlier Mesozoic intrusions. None are alkalic. The K2O abundance is lower (at medium to high K) and most plutons are relatively silicic (>66 wt% SiO2) and calc-alkaline. Two-mica granites are common, as are metaluminous hornblende-biotite-sphene granodiorites. Sr contents are usually less than 800 ppm, except for the calcic and Sr-rich granitoids of the Whipple core complex. Available isotopic data are currently limited to plutons in the Whipple, Chemehuevi, San Gabriel, and Old Woman Mountains and include δ18O values from 7 to 9‰,

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Sri from 0.706 to 0.711 (to 0.719 for the strongly peraluminous granites of the Old Woman Mountains), and εNdt from −10 to −17 (John and Wooden, 1990; Anderson and Cullers, 1990; Barth et al., 1995; Miller et al., 1990). Tested models show a broad range of magma origins for the above complexes, including signiﬁcant derivation from heterogeneous Proterozoic crust, but with some variable input of radiogenic mantle and/or maﬁc crust mixed with a Proterozoic component. The main compositional difference between the Cretaceous plutons is seemingly related to depth of emplacement. While the shallowly emplaced (<12 km) complexes tend to be fundamentally granitic and often leucocratic, deeper Mesozoic plutons in both areas (Tujunga and Mojave) tend to be more maﬁc including metaluminous diorite to quartz diorite and marginally metaluminous to calcic peraluminous tonalite and granodiorite. Most of the shallow plutons appear to have been largely crustal derived. In contrast, the origin of plutonism in the deeper complexes is interpreted to be more variable, including (1) derivation from Proterozoic crust, including granites of the Chemehuevi and Old Woman Mountains (John and Wooden, 1990; Miller et al., 1990), (2) derivation from maﬁc crust with signiﬁcant felsic crustal interaction, including calcic granitoids of the Whipple core complex (Anderson and Cullers, 1990) and (3) mantle-derived, such as the Josephine tonalite of the San Gabriel Mountains (Barth et al., 1995). Directions to Stop 3 Depart stop, return south on Eagle Mountain road to I-10. Continue west on I-10. Note that Chiriaco summit [0618500 3725100] is the last fuel before driving through Joshua Tree National Park. Exit Cottonwood Springs Road [0611140 3725050]. Continue north on Cottonwood Springs Road, past the park entrance [0611170 3726600] to Stop 3 [0610180 3731570].

We have now traveled ~35 km west-southwest, and are now west of the main body of the Middle Jurassic arc visited at Stop 2. As noted there, Mesozoic batholiths of the Mojave Desert and the Tujunga terrane were emplaced in distinct Triassic, Jurassic and Cretaceous pulses. As we move west, we move both structurally deeper and outboard through the arc. At this stop we are structurally deeper, but still at relatively shallow crustal levels. Here we can see a Late Jurassic pluton intruding Proterozoic crystalline basement that underlies supracrustal sequences to the east. Here the pluton is composed of felsic biotite granite cut by maﬁc dikes, characteristic of the relatively outboard, Late Jurassic intrusions of the Jurassic batholithic pulse (Barth et al., 2008). Directions to Stop 4 Depart stop, continue north on Cottonwood Springs Road. Note that the park Visitor Center [0608950 3734665] issues permits to pass through the park, and is also the last water source in the park. Here you can also purchase a comprehensive guide to park geology (Trent and Hazlett, 2002). Note also that there is currently reservable group camping at Cottonwood Campground, located just east of the Visitor Center. Continue north on Cottonwood Springs Road to Stop 4 [0611577 3738280]. Stop 4. Late Jurassic and Late Cretaceous Granite Plutons This stop ﬁnds us in the northern Cottonwood Pass area of the southcentral part of Joshua Tree National Park, between the Hexie Mountains to the west and the Eagle Mountains to the east. Here we have a 360° view of major plutonic rock units that characterize the central part of the park. In this region we are west of the main body of the Middle Jurassic arc, where the PermoTriassic, Late Jurassic and Late Cretaceous arcs overlap. Dated plutons in this area include the Triassic Munsen Canyon pluton (Barth and Wooden, 2006), the Late Jurassic Cottonwood pluton (Barth et al., 2008) and the Late Cretaceous Porcupine Wash and Smoke Tree Wash plutons (Fig. 8). Compositions of plutons in central Joshua Tree National Park illustrate regional secular compositional variation in arc plutons, here expressed in these comparatively homogeneous plutons. Permo-Triassic plutons are typically composed of alkalic monzodiorite to monzonite, and true granite is relatively rare. Rock compositions are characterized by high Sr and low Rb/Sr. Late Jurassic and Late Cretaceous plutons are typically calc-alkalic granodiorite to granite, and have lower Sr and exhibit more linear Rb/Sr covariation (Fig. 9). The Smoke Tree Wash pluton is an excellent example of a discordant, relatively homogeneous pluton that characterizes the central part of the tilted section. It has at least three different phases: the most abundant are coarse and medium-grained phases of k-feldspar, plagioclase, quartz, biotite ± hornblende felsic granodiorite to granite. K-feldspars tend to form larger

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Figure 8. Simpliﬁed geologic map of plutonic units in the Cottonwood Pass area, adapted from Jennings (1967) and Powell (2001a, 2001b).

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Eastern Transverse Ranges and southern Mojave Desert phenocrysts in the medium grained phase. There is also a slightly more maﬁc, medium grained granodiorite, discontinuously exposed near the contact with the host gneisses that is also richer in microgranitoid enclaves. All three units are cut by aplitic to pegmatitic, probably radial dikes that also extend into the host and discordantly cut host fabrics. These granitic dikes are variable in composition some with large k-feldspars, some with two micas and a few with rare garnet. Structural observations clearly separate wall rock deformation from pluton emplacement. The host gneisses comprise both paragneisses and orthogneisses, all strongly foliated with the dominant foliation parallel to the axial planes of isoclinal folds. These folds and foliation are refolded by locally developed more open folds with variable orientations. All of these structures are clearly cut by the pluton and by dikes emanating from the pluton. Magmatic fabrics in the pluton are present, but weak. Orientations are quite variable but there is some suggestion that both the magmatic foliation and locally more intense magmatic lineation have fairly shallow orientations. We don’t yet know if this reﬂects a regional fabric such as that seen in the host or if the shallow orientation reﬂects the presence of a nearby roof. The contact with the host is very discordant, takes numerous decameter-scale steps, and is bordered by xenoliths of host gneiss in the pluton, some of which are clearly rotated and stoped blocks. The contact has a fairly steep dip on the sides of the pluton but then rolls over into a more gently dipping contact in locations that we interpret to be close to the roof of the pluton. Directions to Stop 5 Depart stop, continue north on Cottonwood Springs Road to Stop 5 [0613737 3744989]. Stop 5. Late Cretaceous Granodiorite: Overview of Pinto Basin At this brief stop in the northern part of the broader Cottonwood Pass area, we can examine somewhat more maﬁc granite and granodiorite of Late Cretaceous age, in comparison to the relatively leucocratic granites at the last two stops (Fig. 8). This stop also provides a scenic overview of the topography and geology of the eastern part of the National Park. To the north is Pinto Mountain, composed of Mesoproterozoic sedimentary rocks overlying Paleoproterozoic gneiss and metagranite. The eastern side of this ridge and the lower, dark hills to the right (east) are composed of Jurassic granitic rocks of the Pinto Mountains, similar to those described at Stop 2. In the foreground is Pinto Basin, eroded out of fractured rock along the (transtensional?) Blue Cut fault (Powell, 1981). To the northeast across Pinto Basin are the northern Eagle Mountains, composed of rocks similar to the Pinto Mountains, but the low dark hills at the foot of the Eagle Mountains are alkali basalts of Pliocene age. The distinctly lighter-colored range on the skyline in the background is the Coxcomb Mountains (also seen from Stop

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2), composed of Late Cretaceous granitic rocks intruding the McCoy Mountains Formation. Directions to Stop 6 Depart stop, continue north on Pinto Basin Road to Stop 6 [0591190 3760244]. Stop 6. Late Jurassic White Tank Granite Intruding Paleoproterozoic Gneiss The plutons we will see on this ﬁeld trip, and in all arcs on Earth, show highly variable characteristics. We should thus not be surprised to ﬁnd that there are more than one means by which plutons are emplaced. As we evaluate different emplacement models, we need to keep in mind the following: 1. Plutons grow incrementally or continuously over some duration of time. Thus “emplacement” also occurs over some length of time and may or may not be an episodic process. 2. There is widespread evidence that host is displaced by multiple processes during chamber construction and that these processes vary with depth, distance from the pluton, and time. 3. As younger magma pulses arrive in a growing chamber, the “emplacement problem” may largely involve the displacement of older magmatic pulses rather than the original host. 4. Some “host-rock” transfer processes, such as stoping may remove evidence of earlier processes. 5. The displacement near plutons of regional pre-emplacement markers is an invaluable tool for determining the general direction that host rock is displaced during emplacement. On this trip we will see two general types of plutons (1) roughly circular (in map view) bodies with steep walls that typically cut discordantly across pre-existing structures; and (2) more gently dipping sheeted bodies, often deformed in the magmatic or subsolidus state, that are typically concordant to structures in the host, but on close inspection may have some discordant margins. At this stop we are viewing the southeast margin of the White Tank pluton, an excellent example of the ﬁrst style of pluton. Here the pluton contact has a fairly steep dip (but when mapped in detail is somewhat irregular) and on a regional scale is very discordant to structures in the host gneisses. Locally these structures bend into parallelism with the margin. But only very minor emplacementrelated strain is seen in the contact aureole. Elsewhere we ﬁnd xenoliths of the host gneisses in the pluton, some which may be rafts (unrotated and only slightly displaced xenoliths), and others that are clearly rotated stoped blocks. Host xenoliths are not volumetrically common in the plutons and rapidly decrease in size and number with distance from the pluton margin. One of the exciting new observations from these plutons is that they have geochronologic and/or geochemical data suggesting

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that they are made up of more than one batch of magma. Unfortunately internal contacts between these pulses have not yet been recognized and examined in detail, information needed to fully understand how these bodies grew to their present size and shapes. Even so we can use the shape of the plutons and nature of their contacts with host to consider emplacement mechanisms. It is clear that stoping did affect the ﬁnal characteristics of these contact aureoles and thus we have probably lost information about previously operating emplacement mechanisms. We will discuss the pros and cons of emplacement models such as diapirs, dike fed balloons, sheeted dike complexes, punch laccoliths, and incrementally grown laccoliths and lopoliths at this locality. End of Day 1. DAY 2 Today we will take three moderate hikes to view rock units of the Keys View quadrangle (Fig. 10), situated along the boundary between the middle and lower parts of the tilted crustal section that is the focus of this trip. In this quadrangle, the terrane of Paleoproterozoic gneissic rocks intruded by relatively homogeneous plutons gives way, across a relatively abrupt boundary, to a northwest-trending terrain of layered crystalline rocks lacking discrete mappable plutons such as those we visited on Day 1. This terrain was originally mapped as Precambrian(?) gneissic rocks, and later interpreted as a largely late Mesozoic gneissic complex with pervasive brittle deformation (Powell, 1981). More recent geochronologic work (Needy et al., 2006) shows that this terrain is largely composed of Jurassic and Cretaceous foliated igneous rocks characterized by an intrusive geometry of concordant to slightly discordant meter to decimeter thick sheets, typically dipping moderately north and east. Reconnaissance petrographic work suggests that intrusive sheets record a broad compositional spectrum of intrusions, but that the volumetrically dominant components are tonalite-granodiorite and biotite ± muscovite ± garnet granite sheets. Thermobarometric results suggest that components of this heterogeneous sheeted complex crystallized at 15–22 km, ~5 km deeper than plutons exposed to the east. Similarly layered complexes, recording similar paleodepths, are exposed for an additional 100 km to the northwest (Powell, 1993; Barth et al., 2001b). These results, and the widespread nature of these sheeted rocks, suggest that the sheeted complex represents exposure of the mid-crustal counterpart of the ignimbrites and discordant plutons typical of the upper crust of the Cordilleran continental margin arc at this latitude. Regionally extensive (10s to 100s of km long) magmatic “sheeted complexes” occur in other arcs and/or orogenic belts. These complexes have characteristics that make them very distinct from single sheeted plutons and from adjacent and temporally overlapping magmatic belts such as the following: (1) they often occur in the central core of orogens and/or near prominent tectonic boundaries (2) their dimensions are much greater than an individual pluton (often 100s of km long); (3) magmatic bodies in these complexes ALL have sheet-like or dike-like shapes

Figure 10. Simpliﬁed bedrock geologic map of the Keys View 7.5 min quadrangle. This quadrangle encompasses the boundary between an eastern region of Paleoproterozoic gneiss and discordant plutons, typical of most of the eastern Transverse Ranges, and the structurally lower sheeted complex to the southwest. The Late Cretaceous Palms and Squaw Tank granites are typical discordant plutons, and the Blue granodiorite, Stubbe Springs and Rusty granites, and the Quail Mountain and Bighorn complexes together comprise the sheeted complex in this quadrangle.

versus the more irregular to often elliptical shapes outside them; (4) although magmatism in these zones is typically long-lived (10s to >100 m.y.), they ALWAYS form sheeted bodies, suggesting that emplacement style is consistent over time; (5) they have a dramatic range in composition (gabbros to granites and cumulates to late aplites and pegmatites) over spatially short distances that are not organized into patterns, such as in normally or reversely zoned plutons; and (6) sheeted plutons are usually strongly deformed in the magmatic or solid state indicating the presence of long-lived or recurrent deviatoric stress in the mid crust.

Eastern Transverse Ranges and southern Mojave Desert For example, if we consider Mesozoic magmatism in the North American Cordillera the following prominent sheeted complexes have been recognized: (1) The “Great Tonalite Sill” (really a collection of sills or dikes) in the Coast Plutonic Complex; (2) the Skagit Gneiss Complex in the Cascades crystalline core, Washington; (3) sheeted complexes in the Idaho suture zone; (4) a central sheeted complex in Peninsular Ranges Batholith that extends from southern California to at least as far south as the Sierra San Pedro Martir; and (5) the sheeted complex in the Keys View quadrangle. Such magmatic sheeted zones have been interpreted in a number of ways, such as: (1) the deeper parts of magma plumbing systems that are feeding more elliptically shaped chambers at shallower levels (e.g., Cascades core); (2) magmatism controlled by emplacement into an active fault or faults (e.g., Great Tonalite Sill); (3) magmatism along fundamental lithospheric-scale boundaries that may or may not be faults (e.g., Idaho suture zone and Peninsular Ranges Batholith examples, interpreted to represent a transition from continental to oceanic crust, with or without major faults); (4) dike swarms controlled by a regional stress ﬁeld; (5) sheeted bodies (whether or not dikes) in which emplacement is strongly controlled by preexisting host rock anisotropy; (6) zones of syn-emplacement extension; and (7) combinations of the above. Directions to Stop 7 Depart camp, travel southeast on Quail Springs Road to the intersection with Keys View Road. Continue south on Keys View Road to Stop 7 [0577200 3760650]. Stop 7. Hike West to View Late Cretaceous Palms Granite and Intrusive Contact with Paleoproterozoic Gneiss The grain size, apparent composition, and weak fabrics in the Late Cretaceous Palms pluton do not change their general appearance as the margin is approached, although local variations occur. The intrusive contact is often quite discordant to the metamorphic layering and dominant foliation in the gneisses, but there are some local zones where older structures are rotated into subparallelism with the Palms contact. Some xenoliths (both rafts and stoped blocks) certainly occur along the margin. The intrusive contact changes its orientation from place to place but the overall pattern is that it dips fairly steeply to the northeast. There is one interesting area where the Blue granodiorite has a gently dipping upper intrusive contact with the gneisses, which is then truncated by the fairly steeply dipping Palms contact. This is one location that would indicate that the Palms did intrude more gently dipping sheeted intrusive rocks and gneisses.

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Stop 8. Hike East to View Proterozoic Gneiss and Contact with Late Cretaceous Blue Granodiorite In the Keys View quadrangle (Fig. 10), we can examine a wonderful transition between fairly circular (in map view) plutons that in the ﬁeld appear fairly homogeneous and weakly foliated (Palms granite, Squaw Tank granite), to more elongate plutons with more intense fabrics and clear evidence for multiple internal sheets (e.g., Blue granodiorite, Stubbe Springs granite), to a very heterogeneous sheeted complex showing a range of compositions from gabbro to two mica granites, both Cretaceous and Jurassic ages, and with variable fabric intensities (Quail Mountain complex, Bighorn complex). At this stop, we will hike through the transition from Paleoproterozoic gneisses into the internally sheeted Blue granodiorite. Gneisses here are a mixture of both orthogneisses and paragneisses with local amphibolite. All are intensely deformed—we see clear evidence for early isoclinal folds with the dominant fabric parallel to their axial planes. These early isoclines fold a previous foliation and metamorphic layering and are in turn refolded by more open, upright folds. The regional metamorphic grade was at least amphibolite facies. We also observe local development of andalusite, garnet, and sillimanite near the contact with the Blue granodiorite, but more work is needed to sort out the differences between Paleoproterozoic regional and Mesozoic contact metamorphism. Structures in the gneisses are discordantly intruded along a fairly steeply dipping contact of the Blue granodiorite, although the usual local zones of subparallelism exist. The intrusive contact takes a number of sharp steps and we do ﬁnd xenoliths (some large enough to be mapped) in the granodiorite. Along this eastern margin, the Blue granodiorite tends to be more homogeneous (less sheeted) and has numerous maﬁc clots (small enclaves?) in it. Magmatic fabrics steepen near the margin but are typically fairly gently southwest dipping (foliation) or gently northwestsoutheast–plunging (lineation) in this body. Small compositional and textural changes deﬁne internal sheets in the granodiorite. More dramatic sheeting deﬁned by 2-mica, garnet granites also occurs. Small pods of Stubbe Springs–like granites intrude the granodiorite, particularly as its western margin is approached. Directions to Stop 9 Depart stop, travel west to return to Keys View Road, continue south on Keys View Road to the Keys View parking area, Stop 9 [0575226 3754200]. Stop 9. Hike to View Sheeted Granodiorite and Granite along Western Contact of Late Cretaceous Blue Granodiorite

Directions to Stop 8 Depart stop, continue south on Keys View Road to Lost Horse Mine (dirt) Road, turn left and continue southeast on Lost Horse Mine Road to parking area [0576530 3757780].

The western margin of the Blue granodiorite is often much more cryptic because of the sheets within the granodiorite and younger intrusions along this margin. We will hike along a transect perpendicular to this margin to examine the granodiorite

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and a number of granitic bodies intruding into the granodiorite and end near Keys View where the midcrustal Bighorn complex is exposed. Here the Blue granodiorite is a medium-grained, hornblende biotite granodiorite with a moderately well developed northweststriking, northeast-dipping magmatic foliation and northwest trending lineation. Regionally, the magmatic foliation deﬁnes an open synformal fold pattern with the northwest-southeast–trending fold axis subparallel to the mineral lineation. Locally, zones of subsolidus deformation deﬁne steep zones of southwest-sideup shear. Bodies of 2-mica garnet granite intrude the granodiorite. These bodies are less strongly foliated than the granodiorite and have shapes ranging from sheet-like to blobby. Farther southwest, granodiorites are increasingly intruded by a complex variety of sheet-shaped bodies, locally separated by screens of orthogneiss and paragneiss, and varying in composition from gabbros and hornblende cumulates to 2-mica granites. Sheets vary in thickness from less than a meter to ≥10 m and generally dip moderately to the northeast. Intrusive relationships are complex, but there is a general pattern of more maﬁc being intruded by more

felsic sheets. Many sheets have well-developed magmatic fabrics locally overprinted by weak subsolidus deformation. These magmatic fabrics are more intensely developed in granodiorites and are typically weaker in granites (Brown et al., 2006). End of Day 2 DAY 3 In Days 1 and 2 of this trip, we traveled from east to west through progressively deeper levels of the Cordilleran magmatic arc and its Proterozoic wall rocks. This morning we will visit the northwestern Orocopia Mountains, which in their higher structural levels include crustal rocks of North American afﬁnity similar to those viewed previously. However, the Orocopia Mountains also provide a window into an oceanic complex which was underplated beneath much of southern California and adjoining regions during the Late Cretaceous to early Cenozoic Laramide orogeny. Here known as the Orocopia Schist, these rocks are part of the larger Pelona-Orocopia-Rand Schist (Fig. 11; Haxel et al., 2002; Jacobson, et al., 2007). Emplacement of the schists beneath

Figure 11. Geology of the Orocopia Mountains and vicinity (after Crowell, 1975). Inset shows distribution of Pelona-Orocopia-Rand Schists. Gf—Garlock fault, LA—Los Angeles, SAf—San Andreas fault.

Eastern Transverse Ranges and southern Mojave Desert the arc terrane represents a ﬁrst order tectonic event that involved the stripping away of the lowermost North American crust and the entire thickness of underlying mantle lithosphere. The original subduction thrust responsible for this event is preserved in only a few areas. In most cases it has been replaced by syn-subduction (i.e., Laramide age) and/or middle Cenozoic low-angle normal faults. In the Orocopia Mountains, there is indirect evidence for Laramide exhumation, but the primary contact between schist and arc terrane (here referred to as the “upper plate”) is a major Miocene detachment fault (Orocopia Mountains detachment fault; Robinson and Frost, 1996; Ebert, 2004; Jacobson et al., 2007). Below we provide brief descriptions of the schist and upper plate. Orocopia Schist The schist in the Orocopia Mountains is composed dominantly of metagraywacke interpreted as trench sediment (Grove et al., 2003; Jacobson et al., 2007). Detrital zircons indicate a depositional age no older than ca. 70 Ma and perhaps as young as 62 Ma (Fig. 12; but note that the schist protolith in other

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areas is as old as 90 Ma). The schist also includes various rock types inferred to have been scraped off the Farallon plate. These include a few percent of metabasite with normal to enriched mid-ocean ridge basalt composition and even lesser amounts of Fe-Mn metachert, marble, serpentinite, and talc-actinolite rock (Haxel et al., 2002). Hornblende 40Ar/39Ar ages of 54–50 Ma and muscovite ages as old as ca. 50 Ma (Fig. 12) demonstrate that underplating beneath North American crust and initial cooling and exhumation occurred shortly after deposition of the graywacke protolith in the trench. Biotite ages of 30–20 Ma and Kfeldspar multi-diffusion domain analysis reveal a second, very rapid, period of exhumation at ca. 24–22 Ma (Fig. 12). The latter is attributed to normal-sense slip on the Orocopia Mountains detachment fault. In most of the range, prograde assemblages in the schist belong to the albite-epidote amphibolite faces. However, the upper half of the exposed section shows extensive retrogression to greenschist facies. We believe that this overprint is related to the ﬁrst (early Cenozoic) phase of exhumation. A thin (several m to several 10s of m) mylonite zone right at the top of the schist is

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Time (Ma) Figure 12. Generalized temperature-time paths for Orocopia Schist (left-dipping diagonal lines) and upper-plate rocks (right-dipping diagonal lines) (from Jacobson et al. (2007)). Paths based upon apatite ﬁssion track data (boxes), detrital zircons (black circles), K-feldspar 40Ar/39Ar multi-diffusion domain analysis (snake-like bands), and hornblende (diamonds), muscovite (pentagons), and biotite (hexagons) 40Ar/39Ar bulk closure ages. Schist versus upper-plate ages can be identiﬁed by their proximity to the respective temperature-time bands. The exception is the apatite ﬁssion track age denoted by the white box, which overlaps curves for both the schist and upper plate. This age result is from the schist. Note that divergent paths have been indicated for upper-plate rocks present within the Orocopia Mountains depending upon their proximity to the Orocopia Mountains detachment fault. CMF—Chocolate Mountains fault. See Jacobson et al. (2007) for further explanation.

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considered to be a younger feature associated with the Orocopia Mountains detachment fault. Upper Plate The upper plate includes three main lithologic units: Proterozoic gneiss, 1.2 Ga anorthosite-syenite, and 76 Ma leucogranite (Crowell, 1962, 1975; Crowell and Walker, 1962; Barth et al., 2001a; Jacobson et al., 2007). The gneiss is locally intruded by the anorthosite-syenite, but in general these two units are spatially distinct. In contrast, the gneiss is intimately intruded by the leucogranite, so we combine these two rock types in a single map unit (leucogranite-gneiss) despite their disparate ages. In most areas, the anorthosite-syenite unit sits structurally above the leucogranite-gneiss unit along a low-angle fault referred to as the upper plate detachment fault (Figs. 11 and 13). Additional lowangle faults may also be present within the anorthosite-syenite and leucogranite-gneiss units. The anorthosite-syenite unit includes anorthosite, gabbro, and syenite, as well as compositions (mangerite and jotunite) intermediate between these three end members. Retrograde metamorphism of inferred Proterozoic age is indicated by essentially universal replacement of primary igneous pyroxene by hornblende and biotite, presence of hairline mesoperthite, and local development of blue quartz in the more felsic units. Structurally lower levels of the anorthosite-syenite unit are intruded by dikes and small stocks of the same leucogranite that abundantly intrudes the gneiss. The gneiss is largely quartzofeldspathic but also includes a maﬁc component. Metamorphism was largely to amphibolite faces, but reached granulite facies near intrusive contacts with the anorthosite-syenite. As in the anorthosite-syenite, most granulite assemblages have been retrograded to amphibolite facies, which are commonly associated with blue quartz. The leucogranite associated with the gneiss generally has an aplitic texture. It consists mostly of subequal amounts of quartz, plagioclase, and K-Feldspar with minor biotite that is commonly retrograded to chlorite. The various structural levels of the upper plate show diverse cooling histories (Fig. 12). In all cases, however, the early to middle Cenozoic history is substantially different from that of the Orocopia Schist. It is this disparity which indicates that the Orocopia Mountains detachment fault can be no older than early Miocene. Directions to Stop 10 Depart camp, head southeast on Pinto Basin Road to return to south park entrance station, continue south on Cottonwood Pass Road to southern park boundary and intersection with I-10. Cross I-10 and continue heading south. In 0.4 mi [611074 3724293], turn left onto poorly maintained paved road. In 0.05 mi [611150 3724261], turn right onto dirt road that heads SSW up the alluvial fan toward the Orocopia Mountains. In 0.75 mi, pass power line. In another 0.65 mi, the road enters a canyon in modest hills of

Figure 13. Stop locations, day 3. See Figure 11 for location and explanation of fault hachures. UTM 1 km grid ticks indicated. Abbreviations: an-sy—anorthositesyenite unit, lg-gn—leucogranite-gneiss unit, os—Orocopia Schist.

porphyritic granodiorite. Our limited dating indicates a possible Jurassic age. The road skirts the west side of the canyon. Approximately 0.4 mi after entering the canyon [610368 3721575] take the road that leaves the main wash heading SSW. In 0.5 mi [610174 3720855] pass a small hill of Diligencia Formation. This unit sits unconformably upon the granodiorite we just passed in the canyon. The road approaches a major drainage. At 0.3 mi past the outcrop of Diligencia Formation [610279 3720381] veer to the right off the road traversing the alluvial fan and descend into the wash (it is easy to miss this intersection). Drive up the wash. Pass highly altered outcrops of leucogranite in the low walls of the wash. These are located on the northeast side of the Clemens Well fault, which may be a strike-slip fault with modest to large displacement or a steepened detachment fault related to the Orocopia Mountains detachment fault and upper plate detachment fault (Crowell, 1962, 1975; Powell, 1993; Robinson and Frost, 1996). Continuing up the wash, we cross the Clemens Well fault and pass into the leucogranite-gneiss unit. Stop 10 [610556 3718428]. Start of Traverse, Upper Plate Rocks This is as far as one can drive up the wash (~1.5 mi upstream from our entry point into the wash). Vehicle maneuverability is limited and for the ﬁeld trip we may park a bit below this point. From here we will take a moderate hike up the canyon examining the leucogranite-gneiss and anorthosite-syenite units of the upper plate and the underlying Orocopia Schist. Particular emphasis will be placed on the early Miocene fault contacts between these units. If time permits, we will climb up to the main ridge of the Orocopia Mountains for a view into the Salton Trough. We will examine rocks all along the traverse, with a few key localities broken out as stops.

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We start our hike in the leucogranite-gneiss unit. Note the high degree of brittle deformation and hydrothermal alteration. The leucogranite cuts most of the ductile fabric in the upper plate, but locally shows a modest foliation itself.

Return to vehicles. Depart stop, return to I-10, east on I-10 to Blythe, north on U.S. Highway 95 to Las Vegas.

Stop 11 [610992 3717799]. Upper Plate Detachment Fault

The National Science Foundation (EAR-0106881, 0408730 and 0711119 to APB, EAR-9902788 and EAR-0106123 to CEJ), the National Geographic Society (7214-02), the National Park Service, the Southern California Earthquake Center, Department of Energy, and the Joshua Tree National Park Association provided support for this research. We are grateful to the staff at Joshua Tree National Park for supporting our work in the park, and to Kenneth Brown, Kristin Ebert, Nicole Fohey, Marty Grove, Kristin Hughes, Vali Memeti, Sarah Needy, Emerson Palmer, Jane Pedrick, Geoff Pignotta, Kelly Probst, and Ana Vućić who collaborated with us in the research summarized here. Robert Powell, Paul Stone, Richard Tosdal, and Dee Trent provided guidance and many helpful discussions, and Calvin Miller and Gene Smith provided constructive reviews of the manuscript.

This is a truly outstanding exposure of the contact between the leucogranite-gneiss and anorthosite-syenite units (upper plate detachment fault). The leucogranite-gneiss is exposed at the base of the outcrop. It is overlain along a sharp, low-angle contact by the basal part of the anorthosite-syenite unit which itself is cut by highly sheared dikes of leucogranite. This middle slice is in turn overlain along another low-angle fault by a structurally higher level of the anorthosite-syenite unit without leucogranite. Note the well developed gouge and other indicators of brittle deformation along the various contacts, demonstrating the relatively shallow nature of this fault system. Sense-of-shear indicators are not common but generally show top-NE to top-E sense of movement. Note the steep faults of various orientations within the different structural plates. Stop 12 [610947 3717648]. Orocopia Maﬁc Schists Maﬁc Orocopia Schist with albite porphyroblasts up to 4– 5 mm in diameter. Other major prograde minerals are hornblende and epidote (albite-epidote amphibolite facies). Secondary chlorite is widespread in thin section. Metagraywacke in adjacent outcrops is dominated by quartz, albite, and muscovite with lesser biotite and garnet. Secondary chlorite is present in both units. Stop 13 [611070 3717546]. Orocopia Mountains Detachment Fault Contact between Orocopia Schist and the part of the anorthosite-syenite unit intruded by leucogranite dikes (as seen in the middle plate of Stop 11). Note the mylonite and asymmetric shear bands in the schist, but brittle deformation only in the anorthosite-syenite unit. This contrast in structural behavior conﬁrms the thermochronologic evidence (Fig. 12) that this contact is a normal fault. Stop 14 [611325 3717297]. Orocopia Mountains Detachment Fault The contrast between brittle deformation within the upper plate but mylonitization of the schist is similar to that seen at Stop 13. Here the upper plate is transitional in composition between the leucogranite-gneiss and anorthosite-syenite units (i.e., it includes rocks typical of the anorthosite-syenite suite but exhibits an exceptionally high degree of intrusion by the leucogranite). If time permits, we will walk up this contact to the main Orocopia ridge. Note the mylonitic nature of the schist along the contact, similar to the textures observed at Stop 14.

ACKNOWLEDGMENTS

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Fox, L.K., and Miller, D.M., 1990, Jurassic granitoids and related rocks of the southern Bristol Mountains, southern Providence Mountains, and Colton Hills, Mojave Desert, California, in Anderson, J.L., ed., The Nature and Origin of Cordilleran Magmatism: Geological Society of America Memoir 174, p. 111–133. Gerber, M.E., Miller, C.F., and Wooden, J.L., 1995, Plutonism at the interior margin of the Jurassic magmatic arc, Mojave Desert California, in Miller, D.M., and Busby, C., eds., Jurassic Magmatism and Tectonic of the North American Cordillera: Boulder, Colorado, Geological Society of America Special Paper 299, p. 351–374. Grove, M., Jacobson, C.E., Barth, A.P., and Vućić, A., 2003, Temporal and spatial trends of Late Cretaceous-early Tertiary underplating of Pelona and related schist beneath southern California and southwestern Arizona, in Johnson, S.E., Patterson, S.R., Fletcher, J.M., Girty, G.H., Kimbrough, D.L., and Martin-Barajas, A., eds., Tectonic evolution of northwestern Mexico and southwestern USA: Geological Society of America Special Paper 374, p. 381–406. Hall, D.J., Cohen, L.H., and Schiffman, P., 1988, Hydrothermal alteration associated with the Iron Hat iron skarn deposit, eastern Mojave Desert, San Bernardino County, California: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 81, p. 586–605. Harding, L.E., and Coney, P.J., 1985, The geology of the McCoy Mountains Formation, southeastern California and southwestern Arizona: Geological Society of America Bulletin, v. 96, p. 755–769, doi: 10.1130/00167606(1985)96<755:TGOTMM>2.0.CO;2. Haxel, G.B., Jacobson, C.E., Richard, S.M., Tosdal, R.M., and Grubensky, M.J., 2002, The Orocopia Schist in southwest Arizona: Early Tertiary oceanic rocks trapped or transported far inland, in Barth, A.P., ed., Contributions to crustal evolution of the southwestern United States: Geological Society of America Special Paper 365, p. 99–128. Howard, K.A., John, B.E., and Miller, C.F., 1987, Metamorphic core complexes, Mesozoic ductile thrusts, and Cenozoic detachments: Old Womans-Chemehuevi Mountains transect, California and Arizona, in Davis, G.H. and VandenDolder, E.M., eds., Geologic Diversity of Arizona and its Margins: Excursions to Choice Areas: Arizona Geological Survey Special Paper 5, p. 365–383. Jacobson, C.E., Grove, M., Vućić, A., Pedrick, J.N., and Ebert, K.A., 2007, Exhumation of the Orocopia Schist and associated rocks of southeastern California: Relative roles of erosion, synsubduction tectonic denudation, and middle Cenozoic extension, in Cloos, M., Carlson, W.D., Gilbert, M.C., Liou, J.G., and Sorensen, S.S., eds., Convergent Margin Terranes and Associated Regions: A Tribute to W.G. Ernst: Geological Society of America Special Paper 419, p. 1–37. Jennings, C.W., 1967, Geologic Map of California, Salton Sea Sheet: California Division of Mines and Geology, scale 1:250,000. John, B.E., 1987, Geometry and evolution of a mid-crustal extensional fault system: Chemehuevi Mountains, southeastern California, in Coward, M.P., Dewey, J.F., and Hancock, P.L., eds., Continental Extensional Tectonics: Geological Society [London] Special Publication 28, p. 313–335. John, B.E., and Wooden, J.L., 1990, Petrology and geochemistry of the metaluminous to peraluminous Chemehuevi Mountains plutonic suite, southeastern California, in Anderson, J.L., ed., The Nature and Origin of Cordilleran Magmatism: Geological Society of America Memoir 174, p. 111–133. Mayo, D.P., Anderson, J.L., and Wooden, J.L., 1998, Isotopic constraints on the petrogenesis of Jurassic plutons, southeastern California: International Geology Review, v. 40, p. 421–442. Miller, C.F., 1977, Early alkalic plutonism in the calc-alkaline batholithic belt of California: Geology, v. 5, p. 685–688, doi: 10.1130/00917613(1977)5<685:EAPITC>2.0.CO;2. Miller, C.F., 1978, An early Mesozoic alkalic magmatic belt in western North America, in Howell, D.C., and MacDougall, K.A., eds., Mesozoic Paleogeography of the Western United States: SEPM, Paciﬁc Coast Paleogeography Symposium 2, p. 163–174. Miller, J.S., and Glazner, A.F., 1995, Jurassic plutonism and crustal evolution in the central Mojave Desert, California: Contributions to Mineralogy and Petrology, v. 118, p. 379–395, doi: 10.1007/s004100050021. Miller, C.F., and Wooden, J.L., 1994, Anatexis, hybridization, and the modiﬁcation of ancient crust: Mesozoic plutonism in the Old Woman Mountains area, California: Lithos, v. 32, p. 111–133, doi: 10.1016/0024-4937 (94)90025-6. Miller, D.M., Howard, K.A., and John, B.E., 1982, Preliminary geology of the Bristol Lake region, Mojave Desert, California, in Cooper, J. D., compiler,

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Printed in the USA

The Geological Society of America Field Guide 11 2008

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River James E. Faulds* Nevada Bureau of Mines and Geology, MS 178, University of Nevada, Reno, Nevada 89557, USA Keith A. Howard* U.S. Geological Survey, MS 973, Menlo Park, California 94025, USA Ernest M. Duebendorfer* Department of Geology, Northern Arizona University, Flagstaff, Arizona 86011, USA ABSTRACT In northwest Arizona, the relatively unextended Colorado Plateau gives way abruptly to the highly extended Colorado River extensional corridor within the Basin and Range province along a system of major west-dipping normal faults, including the Grand Wash fault zone and South Virgin–White Hills detachment fault. Large growth-fault basins developed in the hanging walls of these faults. Lowering of base level in the corridor facilitated development of the Colorado River and Grand Canyon. This trip explores stratigraphic constraints on the timing of deformation and paleogeographic evolution of the region. Highlights include growth-fault relations that constrain the timing of structural demarcation between the Colorado Plateau and Basin and Range, major fault zones, synextensional megabreccia deposits, nonmarine carbonate and halite deposits that immediately predate arrival of the Colorado River, and a basalt ﬂow interbedded with Colorado River sediments. Structural and stratigraphic relations indicate that the current physiography of the Colorado Plateau–Basin and Range boundary in northwest Arizona began developing ca. 16 Ma, was essentially established by 13 Ma, and has changed little since ca. 8 Ma. The antiquity and abruptness of this boundary, as well as the stratigraphic record, suggest signiﬁcant headward erosion into the high-standing plateau in middle Miocene time. Thick late Miocene evaporite and lacustrine deposits indicate that a long period of internal drainage followed the onset of extension. The widespread distribution of such deposits may signify, however, a large inﬂux of surface waters and/or groundwater from the Colorado Plateau possibly from a precursor to the Colorado River. Stratigraphic relations bracket arrival of a through-ﬂowing Colorado River between 5.6 and 4.4 Ma. Keywords: Basin and Range, Colorado River, extension, paleogeography, Colorado Plateau *[email protected]; [email protected]; [email protected] Faulds, J.E., Howard, K.A., Duebendorfer, E.M., 2008, Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 119–151, doi: 10.1130/2008.ﬂd011(06). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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In northwest Arizona, the Colorado River crosses an unusually abrupt boundary between the Colorado Plateau and the Basin and Range province (Fig. 1). Essentially ﬂat, relatively unextended strata on the high-standing Colorado Plateau give way to moderately to steeply tilted fault blocks in the Basin and Range province across a system of west-dipping normal faults that includes the Grand Wash fault zone (Lucchitta, 1966, 1979) and South Virgin–White Hills detachment fault (Fig. 2). Unlike other parts of the Colorado Plateau–Basin and Range boundary (e.g., southwest Utah and central Arizona), a broad transition zone is missing in northwest Arizona (Fig. 1). Instead, a 100km-wide region of highly extended crust within the Basin and Range, referred to as the northern Colorado River extensional corridor (Faulds et al., 1990), directly borders the Colorado Plateau on the west. Within the footwall of the Grand Wash fault zone, the western edge of the Colorado Plateau is marked by

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Figure 1. Digital elevation model showing the abrupt western margin of the Colorado Plateau. In contrast to broad transition zones throughout much of Utah and Arizona, the Colorado Plateau gives way abruptly westward to the Basin and Range province in the Lake Mead region of northwestern Arizona. Small white box encompasses the study area in the southern White Hills. The map projection is cylindrical and equidistant with the shape corrected for 37.5° north latitude. Lighting is from the northwest.

the imposing, west-facing fault-line escarpment of the Grand Wash Cliffs, which consist of subhorizontal Paleozoic strata rising ~1.3 km above several east-tilted half grabens in the corridor, including the Grand Wash trough and the Hualapai basin. With respect to the base of the Tertiary section, structural relief across the Grand Wash fault zone commonly exceeds 5 km. Approximately 15–30 km west of the Grand Wash Cliffs, the gently west-dipping South Virgin–White Hills detachment fault dissects the corridor. The South Virgin–White Hills detachment fault is one of the most prominent structures in the northern part of the corridor, as it accommodated as much as 17 km of normal displacement and has many characteristics of classic detachment faults (Duebendorfer and Sharp, 1998; Brady et al., 2000). Thus, the transition between the essentially unextended Colorado Plateau to the highly attenuated Basin and Range occurs across a relatively narrow ~30-km-wide region in northwest Arizona. It is noteworthy that the Colorado River ﬂows transversely across this abrupt strain gradient, having excavated the Grand Canyon within the western part of the Colorado Plateau and traversing orthogonal to the structural grain within the Lake Mead region in the northern part of the extensional corridor (Figs. 2 and 3). The evolution of the Colorado River and Grand Canyon have long fascinated geoscientists, and many models have been proposed for its development (e.g., Powell, 1875, 1895; Blackwelder, 1934; Longwell, 1946; Hunt, 1969; Lucchitta. 1966, 1972, 1979, 1989; Young and Spamer, 2001). Recent work has greatly reﬁned the evolution of the Colorado River, particularly the timing of inception for reaches downstream of the Grand Canyon (Spencer et al., 2001; Faulds et al., 2001b, 2002a; House et al., 2005; Dorsey et al., 2007), models for drainage development (Spencer and Pearthree, 2001; House et al., 2005), and rates of incision within the Grand Canyon (Fenton et al., 2001; Pederson and Karlstrom, 2001; Pederson et al., 2002). However, the relationships between major precursor events and development of the Grand Canyon and lower Colorado River have received less attention. Clearly, the structural and topographic foundering of the Basin and Range province, particularly within the Colorado River extensional corridor, promoted excavation of at least the western part of the Grand Canyon within the high-standing Colorado Plateau. Understanding the spatial and temporal patterns of deformation within the extensional corridor is therefore critical for establishing a physiographic, structural, and temporal framework by which to assess the evolution of the Colorado River and associated drainage systems. On this ﬁeld trip, we will evaluate the timing and nature of Cenozoic structural demarcation between the Colorado Plateau and the Basin and Range province in northwest Arizona (Fig. 3), as chronicled in the stratigraphy of major half grabens in the hanging walls of the Grand Wash and South Virgin–White Hills fault zones. A major goal of the trip is to further elucidate the relations between the stratigraphy and deformational history of these basins with the evolution of the Colorado River.

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Gently dipping normal fault (dashed where concealed) Moderately to steeply dipping normal fault (ball on downthrown side) Strike-slip fault (arrows show relative sense of movement Thrust fault Axial part of Black Mountains accommodation zone, dashed where concealed, showing anticlinal and synclinal segments

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Figure 2. Generalized geologic map of the northern Colorado River extensional corridor. The box surrounds area covered by the ﬁeld trip. Basins: DB—Detrital basin; GB—Gregg basin; GT—Grand Wash trough; HB—Hualapai basin; NWH—northern White Hills basin; LVV—Las Vegas Valley; OA—Overton Arm; SAB—Sacramento basin; SWH—southern White Hills basin. Faults: BG—Blind Goddess fault; CMF— Cerbat Mountains fault; CY—Cyclopic fault; DF—Detrital fault; LBF—Lost Basin Range fault; LMF—Lakeside Mine fault; LMFS—Lake Mead fault system; LVVSZ—Las Vegas Valley shear zone; MS—Mountain Spring fault; NGW—northern Grand Wash fault; SGW—southern Grand Wash fault; SIF—Saddle Island fault; SSW—Salt Spring Wash fault; SVWHD—South Virgin-White Hills detachment fault; WHF— White Hills fault; WF—Wheeler Ridge fault. Major physiographic features: BRP—Basin and Range province; CB—central Black Mountains; CL—Callville Mesa; CP—Colorado Plateau; CM—Cerbat Mountains; DS—Dolan Springs; FM—Frenchman Mountain; GC—Grand Canyon; GM—Garnet Mountain; GWC—Grand Wash Cliffs; HM—Hiller Mountains; HR—Highland Range; LB—Lost Basin Range; LM—Lake Mead; ML—Lake Mohave; MM—Muddy Mountains; MV—Meadview; NB—northern Black Mountains; NE—northern Eldorado Mountains and basin; NM—Newberry Mountains; SB—southern Black Mountains; SE—southern Eldorado Mountains; SP—Snap Point; TB—Temple Bar; TM—Table Mountain Plateau; WHR—Wheeler Ridge. Other structures: AZ—Black Mountains accommodation zone; GBB—Gold Butte block; MPB—Mount Perkins block.

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Figure 3. Map showing ﬁeld trip route for each day and major physiographic features in the northern Colorado River extensional corridor and neighboring parts of the western margin of the Colorado Plateau.

REGIONAL GEOLOGIC SETTING Northwestern Arizona and southern Nevada have occupied a critical position in the western Cordillera since Mesozoic time. In Cretaceous to early Tertiary time, this region marked the northern edge of a large crystalline terrane referred to as the Kingman uplift (Goetz et al., 1975) or Kingman arch (Bohannon, 1984), which was stripped of its Paleozoic and Mesozoic cover by erosion during Paleogene time. Later, during the middle to late Tertiary, the Lake Mead region marked the northern end of the highly extended Colorado River extensional corridor, which

was dominated by ~east-west extension (e.g., Davis et al., 1980; Howard and John, 1987; Davis and Lister 1988; Faulds et al., 1990, 2001a; Campbell and John, 1996). In contrast, the region directly north of Lake Mead was characterized by a complex three-dimensional strain ﬁeld involving strike-slip faulting and north-south shortening, in addition to large-magnitude ~east-west extension (e.g., Weber and Smith, 1987; Anderson and Barnhard, 1993; Anderson et al., 1994; Duebendorfer and Simpson, 1994). In the northern part of the extensional corridor, calc-alkaline magmatism and major east-west extension swept northward in early to middle Miocene time (e.g., Glazner and Bartley,

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary 1984; Gans et al., 1989; Faulds et al., 1999). Magmatism began ca. 18–20 Ma, 1–4 m.y. before the onset of major east-west extension (Faulds et al., 1995, 1999, 2002b; Gans and Bohrson, 1998). Voluminous, early Miocene, generally intermediate composition magmatism was accompanied by little deformation, although mild north-south extension affected some areas (Faulds et al., 2001a). Major east-west extension then battered the region beginning ca. 16–17 Ma in the south and migrating north-northwest to the western Lake Mead region by ca. 13 Ma. Extension ended in most areas by 11–8 Ma. Tertiary extension was accommodated by mainly west-dipping normal faults and east tilting of fault blocks in the Lake Mead area (e.g., Anderson, 1971, 1978; Duebendorfer and Sharp, 1998), whereas east-dipping faults and west-tilted fault blocks dominated to the south in the Lake Mohave region (e.g., Faulds et al., 1995; Fig. 2). The boundary between these oppositely dipping normal fault systems has been referred to as the Black Mountains accommodation zone, which corresponds to a 5–10-km-wide region of intermeshing, oppositely dipping normal faults and abundant extensional folds (Faulds et al., 1990, 2001a, 2002b; Faulds and Varga, 1998; Varga et al., 2004). The east- and west-tilted domains on either side of the accommodation zone are termed the Lake Mead and Whipple domains, respectively (Spencer and Reynolds, 1989). Estimates of extension within the northern part of the corridor range from ~75%–100% (e.g., Faulds et al., 1990; Brady et al., 2000). Thick sections (generally >3 km) of Tertiary volcanic and sedimentary strata rest directly on Proterozoic and late Cretaceous metamorphic and plutonic rock within the bulk of the extensional corridor (Anderson, 1971; Sherrod and Nielson, 1993; Faulds et al., 1995, 2002b; Beard, 1996). Sections are thickest in middle to late Miocene half grabens. The strata typically range in age from early to late Miocene and consist of maﬁc to felsic lavas, ash-ﬂow tuffs, clastic sedimentary rocks, rock avalanche deposits, volcanic breccia, and evaporites. Although preserved to the north, east, and west of the region, Paleozoic and Mesozoic strata are missing from all but the northernmost part of the extensional corridor (i.e., Lake Mead region) owing to signiﬁcant early Tertiary erosion of the Kingman arch. Basement rocks include Paleoproterozoic gneisses, ca. 1.4 Ga granite, Late Cretaceous–early Tertiary peraluminous (two-mica and garnet-bearing) granites, and early to middle Miocene silicic to intermediate plutons and maﬁc to felsic dike swarms. During extension, older units were progressively tilted to steeper dips concurrent with deposition of younger sequences on subhorizontal surfaces. Consequently, many of the basins contain well-developed tilt fanning (i.e., growth-fault sequences), whereby tilts within the synextensional parts of the section progressively decrease upwards. Volcanic units in many half grabens permit precise dating of the timing of extension (e.g., Faulds et al., 1995, 1999, 2002b; Duebendorfer and Sharp, 1998; Gans and Bohrson, 1998; Varga et al., 2004). Tilt fanning indicates that major eastwest extension began 16.7–15.7 Ma and continued at high rates until ca. 13 Ma in a broad region of the corridor extending from the latitude of Kingman, Arizona, on the south to the eastern Lake

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Mead region on the north (Anderson et al., 1972; Beard, 1996; Faulds et al., 1995, 1999, 2002b; Duebendorfer and Sharp, 1998). Major extension then shifted northwestward ca. 13 Ma into the western Lake Mead region, where it continued until ca. 9 Ma (Duebendorfer and Wallin, 1991; Harlan et al., 1998; Castor et al., 2000). Since ca. 8 Ma, the northern Colorado River extensional corridor has experienced only minor tilting and faulting. In contrast to the extensional corridor, the Colorado Plateau has remained tectonically stable through Cenozoic time and is essentially unextended at upper-crustal levels, as evidenced by subhorizontal Paleozoic and Mesozoic strata. Approximately 2 km of Paleozoic and Mesozoic strata caps the Colorado Plateau but is absent in much of the Basin and Range province of central and western Arizona (Peirce, 1985; Lucchitta and Young, 1986). Late Cretaceous marine deposits on the Colorado Plateau (Nations, 1989) suggest nearly 2 km of uplift during Cenozoic time (Parsons and McCarthy, 1995). However, the timing and nature of Colorado Plateau uplift remain controversial, because the transition zone and much of the Basin and Range province are structurally higher than the Colorado Plateau. Thus, both the Basin and Range and Colorado Plateau may have originally been uplifted, perhaps in early Tertiary time, but parts of the Basin and Range province later subsided. The border between the Colorado Plateau and Basin and Range is generally marked by a broad (~50–150 km wide) transition zone containing characteristics of both provinces (Peirce, 1985). In northwestern Arizona, however, an abrupt boundary separates the relatively unextended Colorado Plateau and Basin and Range province (Figs. 1 and 2). Large Miocene half grabens along the eastern margin of the Colorado River extensional corridor chronicle the evolution of this tectonic boundary and also elucidate major events that facilitated development of the Colorado River. A TALE OF THREE BASINS In this section, we describe three major basins along the eastern margin of the Colorado River extensional corridor: the Grand Wash trough, White Hills basin, and Hualapai basin. The Grand Wash trough and White Hills basin are complex, composite easttilted half grabens, whereas the Hualapai basin is a relatively simple east-tilted half graben. Neogene deposits within each of these half grabens have important implications for understanding the tectonic and paleogeographic evolution of this region. 40Ar/39Ar geochronology and geochemical correlations of tephras (tephrochronology) constrain the timing of deformation in both the Grand Wash trough and White Hills basin. The ﬁeld trip will visit the Grand Wash trough on Day 1, northern White Hills basin on Day 2, and southern White Hills and Hualapai basins on Day 3. Grand Wash Trough (Day 1) The Grand Wash trough consists of at least two east-tilted half grabens, which are separated by Wheeler Ridge in the north and the Lost Basin Range in the south (Fig. 2). The eastern half

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Figure 4. View looking north at the Grand Wash trough, Grand Wash Cliffs, and Wheeler Ridge from Airport Point at the north end of Grapevine Mesa. Note the contrast between the subhorizontal strata along the Grand Wash Cliffs and steeply east-dipping strata on Wheeler Ridge.

Colorado Plateau

GC

GWC

Figure 5. View looking east of the Grand Wash trough and western margin of the Colorado Plateau, including Airport Point (AP), Grand Canyon (GC), Grand Wash Cliffs (GWC), Grapevine Mesa (GM), which is capped by the Hualapai Limestone), Grapevine Canyon (GrC), Sandy Point (SP), and Wheeler Ridge (WR). Grapevine Mesa essentially marks the ﬂoor of a late Miocene lake that immediately predates arrival of the Colorado River. Also, note the gentle northeast dip of Paleozoic strata along the Grand Wash Cliffs.

GrC

Grand Wash Trough

GM AP

WR

SP

Lake Mead

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary graben developed in the hanging wall of the west-dipping northern Grand Wash fault and is centered in the Grapevine Wash area. To the west, the Gregg Basin is a relatively narrow east-tilted half graben that lies in the hanging wall of the west-dipping Wheeler Ridge and Lost Basin Range faults. The Wheeler Ridge and Lost Basin Range faults probably represent splays of the Grand Wash fault zone. As Wheeler Ridge dies out to the north of Lake Mead, the two half grabens coalesce to form a large composite basin, at least at exposed levels. Dissection by the Colorado River and its tributaries has produced excellent exposures of the upper part of the Tertiary section in both the Grand Wash trough and Gregg Basin (Figs. 4 and 5). The middle to late Miocene section within the Grand Wash trough (referred to as the rocks of the Grand Wash trough after Bohannon, 1984) includes, in ascending order, at least 250 m of middle to late Miocene fanglomerate, more than 120 m of a sandstone-siltstone facies with locally interbedded gypsum, and as much as 300 m of late Miocene limestone (Figs. 6 and 7A; Longwell, 1936; Lucchitta, 1966; Bohannon, 1984; Wallace,

114° 7'30'' W

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1999; Wallace et al., 2005; Blythe, 2005). The units interﬁnger and thicken eastward toward the deeper parts of the half graben. The conglomerate contains many large boulders (>5 m long) of the 1.4 Ga Gold Butte Granite (e.g., Volborth, 1962; Silver et al., 1977), a megacrystic rapakivi granite derived from the Gold Butte block in the south Virgin Mountains ~6–10 km to the west (Longwell, 1936; Lucchitta, 1966; Lucchitta and Young, 1986; Wallace, 1999; Blythe, 2005). The limestone in the Grand Wash trough is known as the Hualapai Limestone and has been correlated with similar limestone elsewhere in the eastern Lake Mead region (Longwell, 1928, 1936; Lucchitta, 1966). It has been interpreted as either marine (Blair, 1978; Blair and Armstrong, 1979) or nonmarine (Lucchitta, 1966; Faulds et al., 1997; Wallace, 1999). The rocks of the Grand Wash trough are bracketed between ca. 15 and 6 Ma. The older age is based on a 15.3 Ma 40Ar/39Ar date on sanidine from a rhyolite tuff near the base of the section on the west ﬂank of Grapevine Mesa (Faulds et al., 2001b). Younger age constraints include (1) an 8.8 Ma basalt ﬂow (Faulds et al., 2001b) intercalated with alluvial fan deposits shed from the

114° 00' W

Water.

F

WR F

IC

N

60

Quaternary-Pliocene fanglomerate units (Qf,QTf). Colorado River sediments (QTc).

AF

Lake Mead

Basalt of Sandy Point (4.4 Ma).

36° 7'30''N

Paleozoic-clast conglomerate (Tcp).

4

VF

Hualapai Limestone (Thgw, Thgb).

SC F

SC

SP

4 1-5

Lake

4

65

Mead

Sandstone-siltstone units (Tsgw,Tsgb).

AP 1-2

Gypsum.

1-3

Proterozoic-clast conglomerate (Tcg).

3

1-4

Rock avalanche deposits (Tbp, Tbx, Tbpx).

4

LP

4

Proterozoic gneiss. Moderately to steeply dipping normal fault with ball and bar on the hanging wall side.

GWC

50

Wash

GC

Paleozoic strata.

NGWF

WR F

1-9

5

4

5

4

Concealed normal fault with ball and bar on the hanging wall side.

25

4 4

Grap evine

MF

LBF

3

1-8 8

1-6

Meadview

1-7

Meadview, Arizona, city limits. 5

GWC

85

Strike and dip of bedding or layering.

36° 00'N

Intermittent stream. 1-2

Field trip stop.

5 kilometers

Figure 6. Generalized geologic map for the Grand Wash trough, showing ﬁeld trip stops. AF—Airport fault; AP—Airport Point; GC—Grapevine Canyon; GWC—Grand Wash Cliffs; ICF—Iceberg Canyon fault; LBF—Lost Basin Range fault; LP—Lookout Point; MF—Meadview fault; NGWF—northern Grand Wash fault; SCF—Sheep Canyon fault; SCVF—South Cove fault; SP—Sandy Point; WRF—Wheeler Ridge fault.

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A Grand Wash Trough

B

C

D

Tby

Northern White Hills

Southern White Hills

Hualapai Basin

QTcr

QTa

Tr

Th

QTa

QTa

Th

Ts

Ttb

Tsb

Tgy Tb2 Ts

Trc Tha

Tbr

Tcg

Tbs

Tc2

Ts

Tcg Tbr

Tsr Tbr

Tvsy

Ts

Tcg

Trc Tbr

Tvso

Tb1

Ths

Tq Tdt?

Mr

Td

Tbs? m

b t

Tc1

Tca

Xg

Ds Xg Xg

Xg

Figure 7. Generalized stratigraphic columns of major basins near the eastern margin of the northern Colorado River extensional corridor, showing approximate tilts and relative unit thicknesses. (A) Grand Wash trough. QTc—Colorado River sediments; Tby—early Pliocene basalts (e.g., 4.4 Ma Sandy Point basalt; Th, 11–6 Ma Hualapai Limestone; Ts—Sandstone-siltstone facies; Tgy—gypsum; Tcg—conglomerate derived from Gold Butte block; Tbr—megabreccia of Proterozoic or Paleozoic rock; Ths—tuffaceous sedimentary rocks probably correlative with Horse Spring Formation (e.g., Beard, 1996) and containing a 15.3 Ma tuff (Faulds et al., 2001b); Mr—Redwall Limestone; Ds—Sultan Limestone; –Cm—Muav Limestone; –Cb—Bright Angel Formation; –Ct—Tapeats Sandstone; Xg—Paleoproterozoic gneiss. (B) Northern White Hills basin. See map (Fig. 10) for letter symbols. QTa—late Miocene-Quaternary basin-ﬁll sediments; Tr—rounded gravel, sand, and silt; Th—Hualapai Limestone; Tc2—angular, poorly sorted sand and gravel, which has yielded dates from ca. 14.4–10.9 Ma (Blythe, 2005); Tb2—olivine basalt dated 8.3 Ma (Beard et al., 2007); Tb1—lower basaltic ﬂows dated ca. 14.6 Ma (Duebendorfer and Sharp, 1998); Tc1—angular, poorly sorted sandy gravel (granular pattern) and lenses of megabreccia (blocky pattern); Tdt?—ash-ﬂow tuff intercalated with megabreccia and poorly sorted angular conglomerate in Tc1, dated as 15.2 Ma (Duebendorfer and Sharp, 1998), may correlate with the tuff of Mt Davis (e.g., Faulds et al., 2002b); Tbs?—ash-ﬂow tuff that may correlate with the tuff of Bridge Spring (e.g., Faulds et al., 2002b); Xg—Proterozoic gneiss. (C) Southern White Hills basin. Patterns and labels are the same as in Figure 11 except Tbr (megabreccia of Proterozoic rock). (D) Hualapai basin: QTa—Holocene-late Miocene shale, conglomerate, gypsum, and anhydrite; Tha—late Miocene (probably ca. 13–8 Ma) halite and lesser shale and anhydrite; Tcg—locally derived late Miocene fanglomerate; Tvsy—middle Miocene (ca. 16–13 Ma) volcanic and sedimentary rock; Tvso—early to middle Miocene (ca. 20–16 Ma) volcanic and sedimentary rock, possibly resting on a thin section of Cambrian strata; Xg—Proterozoic gneiss, granite, and diabase.

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary Grand Wash Cliffs at Nevershine Mesa (Lucchitta et al., 1986) in the northern part of the trough; (2) a 7.43 ± 0.22 Ma 40Ar/39Ar maximum eruptive age on sanidine from an ash-fall tuff intercalated in the upper part of the Hualapai Limestone at Grapevine Mesa (Wallace et al., 2005); and (3) a 6.0 Ma tephra within the upper Hualapai Limestone in the Temple Bar area to the west of the Grand Wash trough (Spencer et al., 2001). Tilting within the rocks of the Grand Wash trough decreases up-section from ~30° in the lowermost ca. 15 Ma units to <5° in the youngest units (ca. 4.4–6 Ma). The gently to moderately tilted (<30°) lower fanglomerate onlaps moderately to steeply (40–90°) east-tilted Paleozoic strata and Proterozoic gneiss at Wheeler Ridge and in the Lost Basin Range, respectively. The late Tertiary section (ca. 13–4.4 Ma) is generally tilted gently eastward (<10°). The Hualapai Limestone onlaps subhorizontal Paleozoic strata along the Grand Wash Cliffs; its upper part is not cut by the northern Grand Wash fault. In contrast, the Wheeler Ridge fault accommodated ~300 m of offset of the Hualapai Limestone (Fig. 8; Lucchitta, 1966; Wallace et al., 2005), gentle tilting of early Pliocene basalts and Colorado River sediments (Howard et al., 2000), and appears to cut early Pleistocene alluvial fan deposits (Wallace et al., 2005). These relations indicate that major extension in the Grand Wash trough began prior to 15.3 Ma and that the main pulse of extension had ended by ca. 13 Ma. Movement on the northern Grand Wash fault had ceased by 6 Ma; however, activity on the Wheeler Ridge fault continued into at least the early Pliocene and possibly Quaternary time. The rocks of the Grand Wash trough have signiﬁcant paleogeographic implications (Lucchitta, 1966, 1979). For example, the timing of possible uplift of the Colorado Plateau during late Cenozoic time (McKee and McKee, 1972) has been extrapolated

127

from studies of basinal sedimentary deposits within the Grand Wash trough and elsewhere within the lower Colorado River region (Lucchitta, 1979). The westerly provenance of the lower conglomerate unit indicates that no major through-going drainages ﬂowed westward from the Colorado Plateau between ca. 15.3 and 11 Ma. The Hualapai Limestone is also important, as it crops out throughout much of the Lake Mead region proximal to the present course of the Colorado River (Fig. 9). Lucchitta (1966) characterized the limestone as lacustrine based on facies relationships with detrital rocks in the Grand Wash trough. In contrast, Blair (1978), Blair and Armstrong (1979), and Bradbury and Blair (1979) used fossil assemblages, petrography, and δ13C isotopic chemistry to interpret the Hualapai Limestone as marine-estuarine. They further concluded that the Hualapai Limestone marked the northern extent of an ancestral Gulf of California. Because no signiﬁcant late Miocene to recent faulting was documented between the Grand Canyon and lower Colorado River regions, Lucchitta (1979) concluded that the Hualapai Limestone and the presumably marine or estuarine Bouse Formation (e.g., Metzger, 1968; Smith, 1970; Buising, 1990) in the lower Colorado River region were similar in age and deposited at approximately the same elevations (sea level or below). These deposits were therefore used to support 400–900 m of Pliocene-Quaternary uplift of the Colorado River extensional corridor and western part of the Colorado Plateau (Lucchitta, 1979, 1998). This uplift presumably induced rapid down-cutting of the Grand Canyon by the Colorado River since 6 Ma (Lucchitta, 1979, 1989). Recent studies, however, have raised serious questions about these interpretations. For example, Spencer and Patchett (1997) concluded on the basis of 87Sr/86Sr isotopic evidence that carbonates within the late Miocene to Pliocene Bouse Formation

Figure 8. View northeast of Gregg Basin syncline and faulted Hualapai Limestone. This syncline results from easttilting of the Gregg Basin half graben and normal drag along the west-dipping Wheeler Ridge and Lost Basin Range faults. Thus, it is extensional in origin.

128

Faulds et al.

0

10

kilometers

WF

?

sin

FS

eg

g

LM

Mead

D Ba etrit sin al

e ak

TB NGW

FM

?

Ba

?

?

Gr

Z

L

Las Vegas Valle y

LVV S

Overton A rm

dy ud M

. ns t M

Water

Hualapai Limestone (11-6 Ma)

Cenozoic basin fill

Late Miocene gypsum deposits Pabco Mine

Late Miocene limestone near Frenchman Mt (~6 Ma)

Bedrock exposures - generalized

in the lower Colorado River region and the Hualapai Limestone near Temple Bar are lacustrine in origin. Furthermore, evidence from fossil assemblages, petrography, and δ13C–δ18O isotopic geochemistry clearly support a nonmarine (lacustrine) origin for the Hualapai Limestone (Wallace, 1999; Faulds et al., 2001c; Wallace et al., 2005). If not marine or estuarine, the limestone within the Lake Mead and lower Colorado River regions cannot be used as evidence to support (1) late Miocene to recent uplift of the southwestern Colorado Plateau, (2) the northern extent of the ancestral Gulf of California, or (3) rapid down-cutting of the Grand Canyon since 6 Ma. White Hills Basin (Days 2 and 3) A large structural block exposing Proterozoic basement separates much of the Grand Wash trough from basins to the west (Fig. 2). From north to south, this 75-km-long block consists of the Gold Butte block, Hiller Mountains, and White Hills, all forming the footwall to the major west-dipping South Virgin– White Hills detachment fault. The White Hills basin is a large northerly trending group of east-dipping half grabens developed in the hanging wall of the South Virgin–White Hills detachment fault. Although the South Virgin–White Hills detachment has a sinuous trace in the White Hills and is marked by several prominent salients and embayments, a more or less continuous basin

Figure 9. Generalized geologic map showing distribution of late Miocene limestone and gypsum deposits in the Lake Mead area. The limestones are all lacustrine and found proximal to the present course of the Colorado River, suggesting that the source of fresh water for the limestones may have been derived from a drainage network (surface water and/or groundwater) that ultimately evolved into the Colorado River. The presence of late Miocene gypsum in most of these basins beneath the limestone and thick salt deposits in some of the neighboring basins (e.g., Overton Arm, Detrital, and Hualapai basins) further suggests that large playas existed just prior to limestone deposition and possibly adjacent to some of the freshwater lakes during limestone deposition. FM—Frenchman Mountain; LMFS— Lake Mead fault system; LVVSZ—Las Vegas Valley shear zone; NGW—northern Grand Wash fault; TB—Temple Bar; WF—Wheeler Ridge fault.

appears to characterize the hanging wall. However, to date, only the northern and southern parts of this presumably continuous basin have been studied in detail. Because these areas contain notable stratigraphic differences, we describe them separately below as the northern and southern White Hills basins. Northern White Hills Basin (Day 2) The hanging wall of the South Virgin–White Hills detachment fault in the northern White Hills exposes tilted middle to upper Miocene fanglomerate, megabreccia, and volcanic rocks bracketed between 15.2 and ca. 10 Ma, for which upward-decreasing dips indicate deposition during movement on the detachment fault (Duebendorfer and Sharp, 1998; Blythe, 2005; Figs. 7B and 10). They and the South Virgin–White Hills detachment fault are overlain unconformably by upper Miocene fanglomerate, olivine basalt ﬂows, and the Hualapai Limestone, as young as 6.0 Ma. Gypsum and mudstone indicate a depocenter in the western part of the basin, in the Virgin-Detrital trough (Longwell, 1936; Beard et al., 2007). The Hualapai Limestone is the youngest basin ﬁll predating the arrival of and incision by the Colorado River. Two unconformity-bounded sequences comprise the hanging-wall strata of the South Virgin–White Hills detachment fault in Salt Spring Wash, which contains the best exposures of the northern White Hills basin. The lower sequence is juxtaposed directly against highly retrograded crystalline footwall rocks along the

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary

129

114°15′

Quaternary and Pliocene

QT 36°

Tr Miocene

Th Tb2 Tc2 Tb1 Tc1

Figure 10. Generalized map of Salt Spring Wash area. Salt Spring detachment fault is a segment of the South Virgin–White Hills detachment fault. QT—Alluvium (Quaternary and Pliocene); Tr—Ancestral Colorado River deposits (Pliocene); Th—Hualapai Limestone (upper Miocene); Tb2—Olivine basalt (upper Miocene); Tc2— Conglomerate (upper to middle Miocene); Tb1—Pyroxene basaltic andesite (middle Miocene); Tc1— Conglomerate, megabreccia, and tuff (middle Miocene); Yg—Gold Butte Granite (Mesoproterozoic); Xgg—Gneiss and granite (Paleoproterozoic).

Proterozoic Golden Rule peak

35°52.5′

South Virgin–White Hills detachment. This section consists of fanglomerate and megabreccia intercalated with a tuff dated at 15.2 Ma (Duebendorfer and Sharp, 1998). The largest single megabreccia block is >200 m thick, can be traced for >1000 m along strike, and appears to have been derived from the footwall of the South Virgin–White Hills detachment fault. Tilts within this lower sequence range from ~30° to 60°. These rocks are overlain along an angular unconformity by the upper sequence, an interbedded basalt-fanglomerate section that thickens to the north and west. The basal unit of the upper sequence is a basaltic andesite dated 14.6 Ma (Duebendorfer and Sharp, 1998). Similar maﬁc volcanic rocks of this sequence thicken westward in the basin toward the northwestern White Hills, where they have yielded dates of 14.8–14.4 Ma (Cascadden, 1991; Beard et al., 2007). The fanglomerates, which overlie and interﬁnger with the basaltic andesite, fan upward in dip and have been dated between ca. 14.4 and 10.9 Ma on tuff interbeds (Blythe, 2005). A correlative megabreccia deposit in the northern part of the basin is as thick as 120 m, contains blocks as long as 30 m, and dips back toward the Gold Butte block, from which the clasts were derived (Longwell, 1936). The hanging-wall basin thus records debris including rock avalanches shed from the Proterozoic rocks in the footwall block of the South Virgin–White Hills detachment fault while the fault

Yg Xgg

was active. Where exposed, the fault dips from 45° to as little as 16°. The fault cuts sediments estimated as young as ca. 10– 12 Ma, the dips of which roll into the fault (Howard et al., 2003; Blythe, 2005). Based on the dips of enclosing fanglomerates, the South Virgin–White Hills detachment fault ceased activity after deposition of tuffs dated 12.0 and ca. 10.9 Ma (Blythe, 2005) and earlier than a basalt ﬂow dated 8.4 Ma (Beard et al., 2007). Overlying little-deformed, upper Miocene fanglomerate and the Hualapai Limestone lap onto the footwall block, where they ﬁll paleovalleys cut into Proterozoic rocks and bridge across the block to connect with the Gregg Basin on the east (Howard et al., 2003). The interior basins therefore ﬁlled during and after faulting on the South Virgin–White Hills detachment fault until the ﬁlls rose to levels that connected the basins across low parts of the large footwall massif. Some of the resulting interbasin sedimentary connections were later occupied by paths of the Pliocene Colorado River, when this regional stream developed across the region (Howard et al., 2008). Southern White Hills Basin (Day 3) The southern White Hills basin is a large, composite, easttilted half graben in the hanging wall of the Cyclopic fault and northern part of the Cerbat Mountains fault (Fig. 11; Price, 1997;

Tsb

FA T UL

Trc

Kg

II

Trc

Tbs

GREGG BASIN

Trc

Ts

II

Xcg

Xg

Xcg

II

II

Xg

34

Xg

White Hills

Trc

QTa M

Ts

Ts Td

Tsr

26

White Hills anticline

Tbi

Td

Tsr

O

U

S IN TA

Tvc

Ts

N

3-1

Anticline, dashed where concealed

CERBAT MOUNTAINS

Fe

QTa

Xg

35

Strike and dip of layering or bedding

3-5

Location of field trip stop

Qp

Red Lake playa sediments

Ts

Volcanics of the southern White Hills Mainly ~16 Ma basaltic andesite flows

Tvc

Volcaniclastic conglomerate ~20.0 to 18.3 Ma

QTa

Quaternary-Tertiary alluvium

Tsr

Felsic volcanics of southern White Hills, ~16.1 Ma tuffs and rhyolite lavas

Tca

Basal arkosic conglomerate - > ~20 Ma

Ttb

Basalt of Table Mountain Plateau - ~8.7 Ma

Tbi

Basaltic andesite intrusion

Kg

Late Cretaceous peraluminous granite

Tsb

Basalt of Senator Mountain - 9.9 Ma

Td

Volcanics of Dolan Springs - Basaltic andesite lavas

Tbs

Tuff of Bridge Spring - 15.2 Ma

Tq

Volcanics of Dixie Queen mine ~18.5-16.7 Ma dacite-andesite flows

Trc

Conglomerate of Rock Spring ~16.6 to 8.7 Ma

Tbo

Basalt and basaltic andesite flows ~20.0 to 18.0 Ma

A Blind Goddess fault Tq

White Hills fault Tbs

16.0 Ma

Trc

Xg

Tq

Xg Xg

Xcg

1.7 Ga megacrystic granite, hachures denote altered granite

Xg

Paleoproterozoic metamorphic rocks

Tbs Trc Tbs Trc

Ttb

Cerbat Mts. Fault

8.7 Ma

QTa

Trc Trc Ts

Ts Trc

Xg

Td Trc

Tbs

Trc Trc

QTa

QTa

Trc

A' 2000

QTa Tbo

1000

Trc

Trc

0

Trc

Ts Xg

Ttb

Northernmost Cerbat Mts Tbo

Ts

meters

Tq

Trc

QTa

Ts

18.5 Ma

Mountain Spring fault

Cambrian Tapeats Sandstone

Table Mountain Plateau

15.2 Ma

WEST

Gently dipping normal fault; dashed where concealed

20

4 kilometers

2000 feet

ce

Ro

ar Pe

3 miles

Tvc

Dolan Springs

ad

rry

Moderately to steeply dipping normal fault; dashed were concealed

Tvc

Td Ts DETRITAL BASIN

QTa

Xg

Ts

RED LAKE PLAYA

Tbo

AT

Trc

Trc

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Figure 11. (A) Generalized geologic map of the southern White Hills and adjacent areas. (B) Cross section A–A′, shown slightly larger than map scale for ease of viewing but with no vertical exaggeration. The conglomerate of Rock Spring is a thick synextensional unit that thickens eastward in the eastern subbasin of the southern White Hills basin. Note that tilts progressively decrease upward between the upper part of the volcanics of Dolan Springs and basalt of Table Mountain Plateau. Greater tilting within the hanging wall indicates a listric geometry for the Cerbat Mountains and Cyclopic fault zones, which probably merge at depth beneath the eastern subbasin.

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary Faulds et al., 2001b), which collectively mark the southern part of the South Virgin–White Hills detachment fault. As much as 3 km of Miocene volcanic and sedimentary strata accumulated on an erosion surface etched into Paleoproterozoic gneiss and granite, as well as Late Cretaceous peraluminous granite. In ascending order, the stratigraphy in the southern White Hills includes (Fig. 7C) (1) 18.5 to ca. 16 Ma trachydacite-trachyandesite lavas of the volcanics of Dixie Queen Mine (cf. Faulds et al., 1995); (2) basaltic andesite lavas of the volcanics of Dolan Springs; (3) a ca. 16 Ma bimodal sequence of intercalated rhyolite lavas, tuffs, and basaltic andesite ﬂows, referred to as the volcanics of the southern White Hills; (4) the 15.2 Ma tuff of Bridge Spring (cf. Morikawa, 1994; Faulds et al., 1995, 2002b); (5) ca. 16–8.7 Ma synextensional fanglomerates, referred to as the conglomerate of Rock Spring; and (6) 8.7 Ma basalt of Table Mountain Plateau and 9.9 Ma basalt of Senator Mountain. The northerly striking Mountain Spring fault separates the southern White Hills into two distinct lithologic domains, or subbasins (Fig. 11). In the eastern subbasin, thick sections of fanglomerate (conglomerate of Rock Spring) and subordinate volcanic units accumulated in an eastward-thickening wedge bounded by the Cerbat Mountains and Cyclopic faults on the east. In contrast, volcanic rocks dominate the southern and central parts of the western subbasin, which developed in the mutual hanging walls of the Mountain Spring, Cerbat Mountains, and Cyclopic faults. As the volcanic section thins to the north in the northern part of the southern White Hills, however, the distinction between the eastern and western subbasins becomes less conspicuous. The large volcanic component in the western subbasin is more characteristic of half grabens within the bulk of the northern Colorado River extensional corridor (e.g., Anderson, 1971, 1978; Faulds, 1996; Faulds et al., 1995, 2001a, 2002b) and contrasts with the sediment-dominated basins along the eastern margin of the corridor and in the Lake Mead area (e.g., Lucchitta, 1966; Bohannon, 1984; Beard, 1996; Faulds et al., 1997, 2001c). The timing of extension within the southern White Hills is bracketed between ca. 16.7 and 8 Ma. Tilts within the southern White Hills progressively decrease up-section from ~75° in the volcanics of Dixie Queen Mine to ~5° in the basaltic lavas of Table Mountain Plateau (Fig. 12). Major east-west extension probably began ca. 16.7–16.2 Ma during deposition of the lowermost part of the conglomerate of Rock Spring, as evidenced by clasts of Proterozoic gneiss likely derived from surrounding footwall blocks. Although not fully exposed in any single fault block in the southern White Hills, concordant tilts in exposed parts of the lower Miocene section within individual fault blocks indicate little tilting and extension prior to ca. 16.7 Ma. Extension was clearly in full swing, however, during eruption of the ca. 16 Ma volcanics of the southern White Hills, as evidenced by angular unconformities with older units and appreciable tilt fanning (tilts decrease up-section from ~50° to 25°) (Figs. 11 and 12). Tilt fanning suggests that peak extension occurred between ca. 16.5 and 15 Ma. The ﬁnal stages of extension are recorded by minor faulting and gentle tilting of the 8.7 Ma basalts. Although no evidence

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for Quaternary faulting was observed in the southern White Hills, it is possible that some of the faulting temporally correlates with that in the eastern Lake Mead area and may therefore be younger than ca. 6 Ma. The conglomerate of Rock Spring is a particularly important unit in the southern White Hills for elucidating the paleogeographic evolution of the region. It accumulated primarily in the eastern subbasin of the southern White Hills basin east of the Mountain Spring fault (Fig. 11). The conglomerate is bracketed between ca. 16.7 and 8.7 Ma by 40Ar/39Ar ages on underlying and overlying volcanic units. Clasts of Proterozoic gneiss and megacrystic granite, ranging up to ~10 m long, dominate the conglomerates. The matrix-supported texture, subangular clasts, and poor sorting suggest a fanglomerate origin for the conglomerate. The conglomerate forms an eastward-thickening wedge-shaped body in the southern White Hills and is as much as 2.6 km thick proximal to the Cerbat Mountains fault (Fig. 11). The conglomerate of Rock Spring probably correlates with similar conglomerate and rock avalanche deposits in the hanging wall of the northern White Hills basin, where Duebendorfer and Sharp (1998) documented thick middle Miocene rock-avalanche deposits of Proterozoic gneiss derived from the footwall of the detachment. Likely sources for the boulders of megacrystic Proterozoic granite in the conglomerate of Rock Spring include the southern footwall of the Cyclopic Mine fault and Garnet Mountain ~15 km to the east along the western margin of the Colorado Plateau (Figs. 2, 3, and 11). The relatively small body of megacrystic granite in the footwall of the Cyclopic fault, compared to the much larger body at Garnet Mountain, and the abundance of megacrystic granitic clasts within the entire 2.6-km-thick section of conglomerate of Rock Springs suggest that Garnet Mountain was a source for at least some of the detritus. Thus, the conglomerate of Rock Spring may record west-ﬂowing drainages eroding headward into the footwall of the South Virgin–White Hills detachment fault and possibly into the western margin of the Colorado Plateau as early as ca. 16 Ma, essentially at the onset of major east-west extension. Hualapai Basin (Day 3) The Hualapai basin is a gently to moderately east-tilted half graben developed in the hanging wall of the southern Grand Wash fault (Figs. 2 and 13). Despite its proximity to the Colorado River and Grand Canyon, the Hualapai basin remains an internally drained, closed depression. Due to a lack of dissection by tributaries of the Colorado River, synextensional middle to late Miocene strata within the basin is obscured by more recent ﬂatlying sediments (in contrast to the highly dissected Grand Wash trough and White Hills basin). Thus, both the stratigraphy and timing of extension cannot be directly inferred from exposures in the Hualapai basin. Nonetheless, drill-hole and seismic reﬂection data indicate that the Hualapai basin contains a thick (~3.9 km) growth-fault sequence of Miocene sedimentary and volcanic rocks, with tilts decreasing up section from ~25° to 0°. As inferred from analysis of core and seismic reﬂection proﬁles, the

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Figure 12. View to northeast of tilt fanning in southern White Hills basin. East tilts progressively decrease up-section from ~35° in 15.2 Ma tuff of Bridge Spring (Tbs) to ~5° in 8.7 Ma basalts of Table Mountain Plateau (Ttb). Trc—conglomerate of Rock Spring; Tv—volcanic rocks.

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Figure 13. Hualapai basin salt deposit. The Hualapai basin contains a 2.5-km-thick nonmarine salt deposit of probable late Miocene age, as evidenced by well and seismic reﬂection data. (A) 1:1 cross section constrained by a migrated seismic reﬂection proﬁle (from Faulds et al., 1997). Patterns for units are the same as in Figure 7D. SGW, southern Grand Wash fault. (B) Core (9 cm in diameter) showing massive halite. (C) Core showing displacive halite crystals in reddishbrown claystone. This texture indicates deposition induced by groundwater discharge in a continental playa.

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary stratigraphy of the Hualapai basin includes (in ascending order): (1) ~750 m of lower to middle Miocene volcanic and sedimentary rock possibly resting on Cambrian strata and/or Proterozoic gneiss, granite, and diabase; (2) ~335 m of middle Miocene volcanic and sedimentary rock; (3) fanglomerates along the margins that interﬁnger with evaporites in the central part of the basin; (4) up to 2500 m of middle to upper Miocene halite intercalated with minor shale (5%–10%) and anhydrite; and (5) ~600 m of late Miocene–Quaternary shale and lesser amounts of gypsum, anhydrite, and conglomerate (Faulds et al., 1997). The Miocene section in the Hualapai basin (Fig. 7D) is dominated by the 2.5-km-thick sequence of halite, one of the thickest known, nonmarine halite deposits in a continental rift (Faulds et al., 1997). The age of the salt deposit is roughly bracketed between ca. 13 and 8 Ma, because (a) it lies in the upper, more gently tilted part (<10°) of a growth-fault sequence (Fig. 13), and (b) extension in the region peaked ca. 16–13 Ma but continued at lower rates until ca. 8 Ma. Thus, the salt temporally correlates with both the Hualapai Limestone and sandstone-siltstone facies in the Grand Wash trough, as well as with thick fanglomerates in the White Hills basin. The texture and bromine content of the halite and S and O isotopic values of intercalated and capping anhydrite indicate that halite deposition took place in an intracontinental playa that accommodated regional groundwater discharge (Faulds et al., 1997). Thick salt deposits of comparable age are also documented in the Detrital and Overton Arm basins to the west (Mannion, 1963). Thus, thick salt deposits appear to rim the central to eastern parts of the Lake Mead area. The source of the salt remains a mystery but may include chloride-rich Pennsylvanian-Permian redbeds on the Colorado Plateau (e.g., Supai Formation). CENOZOIC PALEOGEOGRAPHIC EVOLUTION In early to middle Tertiary time (Paleocene through Oligocene), prior to the onset of major east-west extension, major streams ﬂowed northeastward from the Basin and Range province onto the Colorado Plateau in northwest Arizona, as evidenced by widespread, southwesterly derived Paleocene-Eocene gravels along the western margin of the Colorado Plateau (Young, 1982). Within the northern Colorado River extensional corridor, these northeasterly ﬂowing streams beveled the northerly trending, north-plunging Kingman arch, stripping away thick sections of Mesozoic and Paleozoic strata and exposing Proterozoic and Late Cretaceous crystalline basement. Because the arch terminated northward in the Lake Mead region, Miocene strata rest on progressively older Mesozoic and Paleozoic strata toward the south in the Lake Mead region (Bohannon, 1984) and directly overlie Proterozoic and Late Cretaceous granite and gneiss throughout most of the extensional corridor. The distribution of the 18.5 Ma Peach Springs Tuff, a regionally extensive ignimbrite (Glazner et al., 1986; Nielson et al., 1990), indicates that major drainages in northwest Arizona continued to ﬂow northeastward onto the Colorado Plateau through early Miocene time (Young and Brennan, 1974). The Peach

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Springs Tuff probably erupted from a source near the southern tip of Nevada (Hillhouse and Wells, 1991) but is found as far east as the Grand Canyon region of the Colorado Plateau (Young and Brennan, 1974; Glazner et al., 1986; Nielson et al. 1990). In addition, sequences of presumably correlative 17–19 Ma maﬁc lavas crop out in both the Garnet Mountain area along the western margin of the Colorado Plateau (Lucchitta and Young, 1986; Wenrich et al., 1996) and in the northern Cerbat Mountains in the eastern part of the extensional corridor (Fig. 2; Faulds et al., 1997, 1999). Thus, in northwest Arizona and southern Nevada, including the western Grand Canyon region, the Basin and Range province was topographically higher than the Colorado Plateau, and no major structural or topographic barriers appear to have separated the two provinces as late as the early Miocene. Middle Miocene east-west extension then fragmented the region into complex arrays of tilted fault blocks (e.g., Anderson et al., 1972; Faulds et al., 1990) and induced the topographic and structural foundering of the extensional corridor relative to the Colorado Plateau. Neogene strata within the corridor chronicle both the evolution of the Colorado Plateau–Basin and Range boundary and development of major drainage systems. Major basins began forming in the hanging walls of major west-dipping normal faults (e.g., Grand Wash fault zone) at or near the western margin of the Colorado Plateau as early as ca. 16.5 Ma. Basin development clearly disrupted the regional northeast-ﬂowing drainage that predominated in early Tertiary time. This is particularly evident in the southern White Hills basin by the thick accumulation of easterly derived fanglomerate (Rock Spring conglomerate) shed from the footwall of the Cyclopic fault and possibly from Garnet Mountain along the western margin of the Colorado Plateau. Gently tilted (<10°) 13–8 Ma strata within the Grand Wash trough and southern White Hills basin suggest that movement on the Grand Wash and related fault zones, as well as development of major hanging-wall half grabens, occurred primarily between ca. 16 and 13 Ma. Furthermore, the distribution of the Snap Point–Nevershine Mesa basalt ﬂow indicates that the Grand Wash Cliffs have changed little since 8.8 Ma. These relations suggest that the current physiography of the Colorado Plateau–Basin and Range boundary in northwest Arizona began developing ca. 16 Ma and was essentially established by 13 Ma (Faulds et al., 2001b). Thus, most of the structural and topographic demarcation between the two provinces had developed by ca. 13 Ma. The antiquity and abruptness of the Colorado Plateau-Basin and Range boundary in this region, as well as the stratigraphic record in the southern White Hills, suggest that signiﬁcant headward erosion into the high-standing plateau began in middle Miocene time. Many deep canyons have since been carved into the Colorado Plateau, the most prominent of which is the Grand Canyon of the Colorado River. It is therefore possible that incipient excavation of the western part of the Grand Canyon by an originally small west-ﬂowing stream also began in the middle Miocene (Faulds et al., 2001b). However, thick middle to upper Miocene evaporite and lacustrine deposits within the Lake Mead and surrounding regions

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(e.g., Lucchitta 1966, 1979; Mannion, 1963; Peirce, 1976; Faulds et al. 1997) indicate that a long period of internal drainage followed the onset of major extension and preceded integration by the Colorado River and full development of the Grand Canyon. Nevertheless, the widespread distribution of such deposits (Fig. 9) may signify a large inﬂux of groundwater that possibly issued from Paleozoic limestone aquifers on the Colorado Plateau (e.g., Hunt, 1969; Huntoon, 1996, 2000; Crossey et al., 2002) and/or surface waters from developing drainage systems, both of which may have been related to subsequent development of the Colorado River (Faulds et al., 2001b, 2001c). It is noteworthy that the Hualapai Limestone is restricted to areas relatively proximal to the present course of the Colorado River, whereas thick salt deposits characterize some of the surrounding basins (e.g., Hualapai, Overton Arm, Detrital). It therefore seems likely that some kind of precursor to the Colorado River supplied fresh water to lakes in the Lake Mead area and that some of these lakes were bordered by large playas. The large inﬂux of fresh water may have also fed large groundwater systems that discharged in some of the surrounding isolated basins (e.g., Hualapai basin), forming unusually thick salt deposits in their wake. The northwestward shift of extension into the western Lake Mead region at ca. 13 Ma is noteworthy due to its possible effect on regional base level (Faulds et al., 2001b). Both the 13–9 Ma event in the western Lake Mead region and localized post–9 Ma deformation accentuated the topographic and structural relief between the extensional corridor and Colorado Plateau that had been largely generated in the 16–13 Ma pulse of extension. The most notable post–9 Ma deformation occurred on the Wheeler Ridge fault in the eastern Lake Mead area (Figs. 2 and 8), which accommodated as much as 300 m of post–6 Ma down-to-thewest displacement of the Hualapai Limestone in the Gregg Basin (Lucchitta, 1966; Howard et al., 2000; Wallace et al., 2005). The post–13 Ma tectonism further lowered base-level in the region and probably rejuvenated down-cutting in the Colorado Plateau, thus facilitating excavation of the Grand Canyon and ultimate development of the through-ﬂowing Colorado River. Arrival of Colorado River Several basins in southern Nevada and northwest Arizona record a transition from lacustrine to ﬂuvial deposition between ca. 6 and 4 Ma that marks the arrival of the Colorado River and presumably main phase of excavation of the Grand Canyon (Faulds et al., 2001c, 2002a; House et al., 2005). As previously discussed, the late Miocene landscape in the Lake Mead region was dominated by a series of lakes or wetlands within which the Hualapai Limestone and temporally correlative lacustrine deposits accumulated. The lakes stretched from the mouth of the Grand Canyon in the Grand Wash trough, through the Temple Bar–northern Detrital basin, and westward to the Boulder Basin in the western Lake Mead area. The ca. 11–6 Ma Hualapai Limestone is the youngest deposit formed prior to integration of the Lake Mead region into a through-ﬂowing Colorado River. In the

Gregg Basin, the 4.4 Ma basalt ﬂow at Sandy Point is intercalated in Colorado River gravels. Farther west, a lacustrine limestone just north of Frenchman Mountain rests on the ca. 5.6 Ma Wolverine Creek tephra and interﬁngers eastward with a gypsum deposit that extends to near the modern Colorado River (Castor and Faulds, 2001). These relations bracket Colorado River inception in the Lake Mead region between ca. 5.6 and 4.4 Ma. To the south in the Lake Mohave and Laughlin–Bullhead City area, alluvial fans dominated the late Neogene landscape until ca. 5.6 Ma when small lakes formed (House et al., 2005). Near Bullhead City, the 5.6 Ma Wolverine Creek tephra lies directly below a thin limestone. Nearby, younger Colorado River gravels contain the 3.6–4.2 Ma “lower Nomlaki” tephra. Thus, inception of the Colorado River in the Lake Mohave area is bracketed between ca. 5.6 and 4.2 Ma (House et al., 2005), similar to that in the Lake Mead region. Farther south in the Salton Trough region of southern California, the ﬁrst appearance of Colorado River sand occurred ca. 5.3 Ma (Dorsey et al., 2005, 2007). These relations indicate relatively rapid, regional inception of the lower Colorado River in the early Pliocene throughout northwest Arizona, southern Nevada, and southern California. The chemistry of late Neogene lakes in this region may hold important clues to the mechanisms of inception and is currently under investigation by Roskowski et al. (2007). DAY 1. LAS VEGAS TO GRAND WASH TROUGH Introduction Figure 3 shows an overview of the trip route for each day. The trip focuses on the northwest corner of Arizona between Lake Mead on the north and Kingman, Arizona, on the south. On Day 1, we will drive southeast from Las Vegas directly to the Grand Wash trough in the eastern Lake Mead area, crossing the Colorado River at Hoover Dam on the way. Stratigraphic and structural relations within the Grand Wash trough and Gregg Basin are the focus of Day 1. The Grand Wash trough is a composite, east-tilted half graben in the hanging wall of the Grand Wash fault zone. Flat-lying Paleozoic strata on the Grand Wash Cliffs bound the trough on the east. After overview stops along U.S. Highway 93 approximately 13 miles south of Hoover Dam and on Grapevine Mesa at Airport Point in the Grand Wash trough, subsequent stops will visit (1) a thick section of conglomerate shed from the steeply tilted Gold Butte block to the west, (2) the Wheeler Ridge normal fault, which is part of the Grand Wash fault system and accommodated tilting of post–5 Ma Colorado River sediments, (3) the 4.4 Ma Sandy Point basalt that interﬁngers with Colorado River sediments, and (5) the ca. 11–6 Ma Hualapai Limestone, which immediately predates arrival of the Colorado River. The stratigraphy within the Grand Wash trough records the structural and paleogeographic evolution of the region and will therefore facilitate signiﬁcant discussion of the evolution of both the Colorado Plateau–Basin and Range boundary and the Colorado River.

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary En Route Discussion As we leave Las Vegas, we will travel southeast on U.S. Highway 93/95 and continue along Highway 93 through Boulder City. Highway 93 crosses the Colorado River at Hoover Dam at the head of Black Canyon. Black Canyon contains a thick (>2 km) sequence of moderately to steeply east-tilted early to middle Miocene volcanic rocks (mainly basaltic andesite lavas), which rest directly on Paleoproterozoic basement (Anderson, 1978). Most of the major road-cuts near Hoover Dam, however, are in the 13.9 Ma tuff of Hoover Dam, a dacitic, poorly to moderately welded, ash-ﬂow tuff (Mills, 1994). As we skirt Black Canyon along Highway 93, we traverse through multiple road-cuts of incised late MiocenePliocene alluvial fan sediments, which include a few large megabreccia deposits of Proterozoic gneiss and are locally capped by 4.3–5.9 Ma basalt lavas (Feuerbach et al., 1993). These deposits are essentially untilted and overlie east-tilted strata as young as 12.6– 11 Ma. Thus, they record a largely post-extensional ca. 11–5 Ma episode of erosion of the Wilson Ridge crystalline terrane and subsequent dissection by the Colorado River and its tributaries. Directions to Stop 1-1 Proceed from Las Vegas to Hoover Dam. From Hoover Dam, drive ~12.95 miles south on U.S. Highway 93, then pull off to right at overlook. Stop 1-1. Willow Beach Overlook The Willow Beach overlook provides sweeping views to the west of Black Canyon along the Colorado River and much of the northern Eldorado and northern Black Mountains. The classic early studies of Anderson (1971, 1978) and Anderson et al. (1972) of large-magnitude Miocene extension were conducted in this region. The Colorado River is below, and the Eldorado Mountains lie to the west of the Colorado River. The high ridge to our east is Wilson Ridge, a large horst block dominated by the ca. 13 Ma Wilson Ridge pluton (Larsen and Smith, 1990) in the north and Paleoproterozoic gneisses in the south. Most of the region within view is part of the Lake Mead extensional domain (Spencer and Reynolds, 1989), which is dominated by steeply east-tilted fault blocks bounded by gently to steeply west-dipping normal faults. Malpais Flattop Mesa is the prominent basalt capped mesa to the southwest. 40Ar/39Ar dating of variably tilted volcanic units within an east-tilted half graben along the west ﬂank of Malpais Flattop Mesa brackets major deformation in this area between ca. 15.9 and 11 Ma (Faulds, 1999; Faulds et al., 1999). The gently (~5°) east-tilted basalt lavas that cap Malpais Flattop Mesa have yielded 40Ar/39Ar ages of 11.3–11.6 Ma. Thus, the capping basalts on Malpais Flattop Mesa are appreciably older than the 4.3–5.9 Ma basalts that overlie the fanglomerates on the west ﬂank of Wilson Ridge. En Route Discussion. About 3.5 miles south of the Willow Beach overlook, we travel through Housholder Pass, where the highway descends into Detrital Valley and expands to four lanes.

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Note the contrast between the subdued, relatively undissected topography of Detrital Valley and the highly dissected Black Canyon area. Seismic reﬂection proﬁles indicate that Detrital Valley is bordered on the west by a major steeply east-dipping normal fault (Fig. 2). These proﬁles also suggest that highly reﬂective middle crust indicative of synextensional ductile fabrics (e.g., McCarthy et al., 1991) lies at a relatively shallow depth (~5–6 km) beneath this area (Faulds, 1999). Drill holes show that the Detrital basin contains several hundred meters of late Miocene halite and gypsum. As we continue south on Highway 93, Mount Perkins in the central Black Mountains comes into view on the west. Mount Perkins is a steeply west-tilted fault block bounded on its east ﬂank by a major gently east-dipping normal fault. The Mount Perkins block exposes an ~9 km thick crustal section, including parts of a large stratovolcano and rhyolite dome complex (Faulds et al., 1995). The west-tilted Mount Perkins block lies within the Whipple domain (Spencer and Reynolds, 1989), which is dominated by west-tilted fault blocks bounded by east-dipping normal faults. To our west, the Black Mountains accommodation zone separates the east-tilted Lake Mead domain from the west-tilted Whipple domain (Faulds et al., 1990, 2001a). After we turn east onto the Pearce Ferry Road and pass through Dolan Springs, the highway passes through a broad saddle between the southern White Hills on the north and northern Cerbat Mountains to the south (Fig. 3). A large composite east-tilted half graben comprises the southern White Hills and contains ~3 km of middle to upper Miocene synextensional strata (Price, 1997; Price and Faulds, 1999; Faulds et al., 2001b). Table Mountain Plateau is the prominent basalt-capped mesa to the north and consists of gently east-tilted 8.7 Ma basalt ﬂows. More steeply east-tilted strata form hogback ridges directly west of the Table Mountain Plateau. The southern White Hills basin lies in the hanging wall of the Cyclopic-Cerbat Mountains fault (southern part of the South Virgin–White Hills detachment fault) and will be the focus of Day 3 of the ﬁeld trip. After passing through the saddle, the Pearce Ferry Road gently descends into the north end of the Hualapai basin before climbing onto Grapevine Mesa in the southernmost part of the Grand Wash trough. As we ascend onto Grapevine Mesa, a large westprotruding promontory of the Colorado Plateau, Garnet Mountain, is on the right, and the south end of the Lost Basin Range lies to the left. Once atop Grapevine Mesa, we travel northward through a thick Joshua tree forest and gradually descend toward Airport Point at the north end of Grapevine Mesa. Directly east of Grapevine Mesa lies the imposing fault-line escarpment of the Grand Wash Cliffs, which consist of very gently (<~3°) northeast-tilted Paleozoic strata in the footwall of the northern Grand Wash fault (Figs. 1, 4, and 5). The Grand Wash Cliffs mark the abrupt western margin of the Colorado Plateau in this region. Directions to Stop 1-2 From the Willow Beach overlook (Stop 1-1), continue south on U.S. Highway 93 for 28.8 miles, then turn left onto the Dolan Springs–Pearce Ferry Road. Continue on the Pearce Ferry Road for

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43.8 miles, traversing through the town of Dolan Springs (~6.0 miles from the turnoff), north end of the Hualapai basin, and Grapevine Mesa along the way. At 43.8 miles past the turnoff from Highway 93, turn right onto a dirt road and travel 2.8 miles to Airport Point. Stop 1-2. Airport Point, Grapevine Mesa, Grand Wash Trough Airport Point provides a panoramic view of the abrupt transition between the Colorado Plateau and Basin and Range province (Fig. 4), including the Grand Wash Cliffs and mouth of Grand Canyon to the east, steeply tilted fault blocks and Miocene synextensional strata within the Grand Wash trough to the north, and the steeply tilted Gold Butte block to the west. Airport Point lies in the central part of the Grand Wash trough at the north end of Grapevine Mesa, which is a broad plateau capped primarily by the ca. 11–6 Ma Hualapai Limestone. It is noteworthy that the upper relatively ﬂat surface of Grapevine Mesa, developed on the Hualapai Limestone, essentially marks the ﬂoor of a lake that immediately predates arrival of the Colorado River. Early studies in this area include those of Longwell (1936, 1946), Lucchitta (1966, 1972, 1979), and Bohannon (1984). Lucchitta’s seminal work, which integrated the stratigraphic and structural framework with the evolution of the Colorado River, produced many concepts that have withstood decades of additional research. More recent work in this area has produced more detailed maps (Howard et al., 2003; Wallace et al., 2005; Brady et al., 2008), constrained the ages of critical units and events with 40 Ar/39Ar geochronology and geochemical correlations of tephras (i.e., tephrochronology; Faulds et al., 2001b, 2001c; Wallace et al., 2005), reﬁned the depositional environment of the Hualapai Limestone (Wallace, 1999; Faulds et al., 2001c) and other clastic units (Blythe, 2005), and elucidated the structural and paleogeographic evolution of the region (Fryxell et al., 1992; Fitzgerald et al., 1991, 2003; Brady, 1998; Brady et al., 2000; Reiners et al., 2000; Howard et al., 2000, 2008; Faulds et al., 2001b, 2001c). To our west, the steeply east-tilted Gold Butte block is the northern and widest part of the footwall to the South Virgin–White Hills detachment fault (Figs. 2 and 3). This block has been interpreted as a tilted oblique section through the upper 15–18 km of the crust (Wernicke and Axen, 1988; Fryxell et al., 1992). A series of thermochronologic proﬁles across the Proterozoic gneiss and granite in the block, using a variety of minerals and techniques, indicate rapid cooling from tectonic unrooﬁng ca. 15 Ma (Fitzgerald et al., 1991; Reiners et al., 2000; Reiners, 2002). The Mesoproterozoic Gold Butte Granite in the block was long ago recognized as the source of large boulders in Miocene debris that spread into the Grand Wash Trough as far as the present-day mouth of the Grand Canyon (Longwell, 1936; Lucchitta, 1966). Debris in middle Miocene deposits at Frenchman Mountain since transported far to the west, near Las Vegas, has also been attributed to sources in the Gold Butte block (Longwell, 1974). Longwell (1936) inferred from the coarseness and wide distribution of sediments derived from the block that the range towered above its surroundings.

To the east of Airport Point, the western margin of the Colorado Plateau is marked by the imposing fault-line escarpment of the Grand Wash Cliffs, which rise ~1.3 km above the Grand Wash trough. Paleozoic strata on the Grand Wash Cliffs are tilted very gently (generally <3° and commonly <1°) to the northeast. This northeast-tilting is inherited from regional uplift to the west and southwest during Laramide time and, in this region, essentially marks the north to northeast ﬂanks of the Kingman arch (e.g., Bohannon, 1984; Peirce, 1985). As a result of this gentle northeast tilting, a series of northwest-trending erosional escarpments developed during early Tertiary time in Paleozoic and Mesozoic strata across the southwestern part of the Colorado Plateau. One prominent escarpment in this region formed at the base of resistant formations of Permian limestone (Toroweap and Kaibab Formations) and is still very conspicuous in the present physiography of the Grand Canyon region. The west-dipping northern Grand Wash fault lies near the base of the Grand Wash Cliffs but is onlapped by late Miocene strata within the Grand Wash trough (Lucchitta, 1966). Fault blocks in the hanging wall of the northern Grand Wash fault, such as Wheeler Ridge to our north and the Gold Butte block to our west, are tilted steeply eastward (>60°), suggesting that the northern Grand Wash fault has a listric geometry. The contrast between the nearly ﬂat-lying Paleozoic strata along the Grand Wash Cliffs and the steeply east-tilted Paleozoic strata on Wheeler Ridge (just 10 km west of the Grand Wash Cliffs), which can be easily seen to the north of Airport Point (Fig. 4), epitomize the abrupt transition between the Colorado Plateau and Basin and Range province in this region. It is noteworthy that the aforementioned, prominent northwest-trending early Tertiary escarpment at the base of the resistant Permian limestones on the Colorado Plateau is also preserved on Wheeler Ridge but has been tilted steeply eastward along with the rest of the fault block (Lucchitta and Young, 1986). Progressive erosional beveling toward the south removed the north- to northeast-tilted Paleozoic strata from most of the northern Colorado River extensional corridor, where Miocene strata generally rest directly on Proterozoic basement. Signiﬁcant displacement along the listric, northern Grand Wash fault zone generated the east-tilted composite half of the Grand Wash trough, which ﬁlled primarily with middle to upper Miocene sedimentary deposits. Tilt fanning within these sedimentary rocks indicates that the main pulse of extension began prior to ca. 15.3 Ma and ended by ca. 13 Ma; however, the Hualapai Limestone and an overlying 4.4 Ma basalt intercalated in Colorado River sediments are locally tilted and faulted, demonstrating that some extension continued into Pliocene and possibly Quaternary time. It is noteworthy that an 8.8 Ma basalt lava ﬂowed down the Grand Wash Cliffs from the Snap Point area to Nevershine Mesa (Figs. 2 and 3), indicating that the physiography of the Grand Wash Cliffs has changed little since late Miocene time (Faulds et al., 2001b). On subsequent stops, we will view critical parts of the stratigraphic section within the Grand Wash trough and discuss how parts of this section reﬂect major deformational events and may or may not be related to development of the Colorado River.

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary Directions to Stop 1-3 From Stop 1-2, return to the Pearce Ferry Road and turn right (north). The road descends through a thin veneer of Hualapai Limestone, the underlying sandstone-siltstone facies, and then into a thick section of coarse conglomerate. After 1.5 miles on the Pearce Ferry Road, turn left onto the South Cove Road and travel 0.9 miles to Stop 1-3 (pull off on right shoulder). Stop 1-3. Roadcuts in Fanglomerate Facies Roadcuts here expose a thick sequence of massive, gently east-dipping, generally matrix-supported conglomerate. The conglomerate contains many large boulders (some >5 m long) of the 1.4 Ga Gold Butte Granite, a megacrystic rapakivi granite derived from the Gold Butte block in the south Virgin Mountains ~6–10 km to the west (Longwell, 1936; Lucchitta, 1966; Lucchitta and Young, 1986; Wallace, 1999; Blythe, 2005). Easterly dips in the 250-m-thick section of conglomerate on the west ﬂank of Grapevine Mesa decrease upward from ~25°to ~5°. Rock avalanche megabreccia deposits are locally intercalated in the conglomerate, especially in the lower part of the section. The conglomerate is roughly bracketed between ca. 15.3 and 11 Ma. The poor sorting, angularity and size of clasts, both matrixand clast-supported beds, and intercalated lenses of sandstone suggest that the conglomerate originated as debris-ﬂow and sheetﬂood deposits on alluvial fans. The conglomerate was deposited in large alluvial fan complexes shed eastward from the Gold Butte block and in some cases ﬁlled paleocanyons cut into east-tilted Paleozoic strata (Lucchitta and Young, 1986). These fanglomerates reﬂect relative uplift of the Gold Butte block and subsidence of the Grand Wash trough, both induced by signiﬁcant east-tilting in the hanging wall of the northern Grand Wash fault zone and possibly some isostatic rebound of the Gold Butte block, which lies in the footwall of the South Virgin–White Hills detachment fault. Their westerly provenance indicates that no major through-going drainages ﬂowed westward from the Colorado Plateau during this interval (Lucchitta, 1966). Directions to Stop 1-4 Continue to the west on the South Cove Road, descending through the massive conglomerate and passing the Miocene unconformity on the right, where the conglomerate rests in angular unconformity on Paleozoic strata. After 1.2 miles, pull off to right into large parking area. Stop 1-4. Wheeler Ridge Fault and Gregg Basin We will traverse up the hill to the southeast of the South Cove Road to the Wheeler Ridge fault. The fault trace is marked by steeply (5°–70°) west-dipping ﬂat-irons of the Hualapai Limestone dragged along the fault and juxtaposition of the limestone against the conglomerate that contains clasts of the Gold Butte Granite. The Wheeler Ridge fault bounds the Gregg Basin on the east. The Gregg Basin is a narrow east-tilted half graben that

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merges northward with the Grand Wash trough. Our traverse provides an excellent southward view of the Gregg Basin, including a hanging-wall syncline induced by east-tilting of the half graben and west-facing drag along the Wheeler Ridge fault (Fig. 8). The Wheeler Ridge fault accommodated ~300 m of down-to-the-west normal displacement of the ca. 11–6 Ma Hualapai Limestone (Lucchitta and Young, 1986; Wallace et al., 2005), and its hanging wall exhibits an east-facing rollover anticline in the Gregg Basin, affecting Colorado River sediments intercalated with a 4.4 Ma basalt ﬂow at Sandy Point (Howard et al., 2000). Directions to Stop 1-5 Continue 1.5 miles downhill to the west on the South Cove Road, then veer right and travel 0.3 miles to the picnic area for lunch. Stop 1-5. Sandy Point Viewpoint A short stroll to the west from the picnic area provides a view to the north of the peninsula of Sandy Point, where a 4.41 ± 0.03 Ma basalt ﬂow (Faulds et al., 2001c, from M. Kunk, 1998, written commun.) is intercalated in Colorado River sediments (Fig. 14). This basalt ﬂow provides an important age constraint on the inception of the Colorado River in the Lake Mead area. The Hualapai Limestone, which predates arrival of the Colorado River, is as young as ca. 6 Ma (Spencer et al., 2001). Thus, inception of the Colorado River in the eastern Lake Mead area is constrained between ca. 6 and 4.4 Ma. The sands and rounded gravels in the picnic area are Colorado River sediments and partly correlate with the Chemehuevi Formation of Longwell (1936, 1946). Across Lake Mead on the west side of Gregg Basin, the prominent Jumbo Pass wind gap between the Virgin Mountains (to the north) and the Hiller Mountains (to the south) contains rounded river pebbles recording an abandoned high-level paleovalley of the Colorado River across the Gold Butte–Hiller Mountains–White Hills massif (Howard et al., 2003). The Jumbo Pass area may have been a major late Miocene pathway for debris transported from the Gold Butte block toward the Grand Wash trough, as suggested by coarse-grained upper Miocene fanglomerate exposures near and east of the Pass. Directions to Stop 1-6 Return to Pearce Ferry Road by driving 3.8 miles east on South Cove Road. Turn right (south) toward Meadview onto Pearce Ferry Road and continue south 8.9 miles, then pull off to right to optional stop at Gregg Basin overlook. Stop 1-6. Gregg Basin Overlook (Optional) This overlook lies along the western rim of Grapevine Mesa and therefore provides excellent views to the west of Gregg Basin, the Gold Butte block, Virgin River Canyon along the Colorado River (ﬁlled by Lake Mead), and in the distance the Temple Bar–Detrital basin area. The Virgin River Canyon is one of

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three places where the Colorado River cut through the large Gold Butte tilt block. Whether the Gregg Basin–Grand Wash trough and Temple Bar–Detrital basin were connected prior to arrival of the Colorado River will be discussed on Day 2. In the foreground on the east side of Gregg Basin are two exposures of Proterozoic rock separated by the steeply dipping, poorly exposed Meadview fault, which is marked by a 50–100m-thick damage zone (Wallace et al., 2005; Swaney, 2005). The easternmost Proterozoic exposure lies beneath the Cambrian nonconformity and has yielded a relatively old apatite ﬁssiontrack age (ca. 125 Ma; Fitzgerald et al., 2003). The exposure west of the Meadview fault, at the northernmost tip of the Lost Basin Range, yielded a relatively young apatite ﬁssion-track age of ca. 15 Ma (Fitzgerald et al., 2003). Based on the different cooling and inferred exhumation histories in these adjacent blocks, the Meadview fault appears to be a relatively underappreciated signiﬁcant fault in this region. The Colorado River (and now Lake Mead) exits Gregg Basin westward through narrow Virgin Canyon across the Proterozoic rocks of the footwall block of the South Virgin–White Hills detachment fault. Abandoned paleovalleys containing Colorado River pebbles at Jumbo Pass and also south of Virgin Canyon at Spring Canyon mark former courses of the river across the block. The Hualapai Limestone bridges across the block south of Virgin Canyon and demonstrates that in the late Miocene the Gregg basin merged westward with the northern White Hills basin.

and interbedded gypsum appear to rule out any rigorous and/or constant through-going drainage, at least in exposed areas. The fanglomerate and sandstone-siltstone facies are probably related, but the grayish-brown matrix in the Proterozoic-clast fanglomerate contrasts with the reddish sandstone and siltstone. This suggests that at least two different source areas fed sediment into the Grand Wash trough. The Proterozoic-clast conglomerate was probably derived primarily from the crystalline terranes of the south Virgin Mountains and possibly Lost Basin range, whereas the sandstone-siltstone facies may have been largely derived from nonresistant Pennsylvanian-Permian redbeds (e.g., Hermit and Supai Formations) on the Colorado Plateau and/or to the north of the Grand Wash trough. The exposed part of the sandstone-siltstone facies within this area is bracketed between ca. 13 and 11 Ma. Nonwelded tuffs in the lower part of the exposed section in the Pearce Ferry area have yielded ﬁssion-track ages ranging from 10.8 ± 0.8 to 11.6 ± 1.2 Ma (Bohannon, 1984) and a maximum 40Ar/ 39Ar age on sanidine of 13.11 ± 0.08 Ma, whereas a tephra in the upper part of the section near Airport Point geochemically correlates with a 10.94 ± 0.03 Ma tuff (Wallace et al., 2005).

Directions to Stop 1-7 Continue south on Pearce Ferry Road 0.6 miles, then turn left onto Meadview Boulevard and continue east past stop sign at Meadview Market. After 2.0 miles on Meadview Boulevard, turn right onto Shore Avenue. Travel south 0.1 miles on Shore Avenue, then turn left onto unmarked dirt road (Grapevine Wash Road). Continue 1.1 miles on Grapevine Wash Road to Stop 1-7. Note that as the Grapevine Wash Road is sandy and locally rough, high clearance and four-wheel-drive are advised.

Stop 1-8. Fanning dips in Hualapai Limestone

Stop 1-7. Sandstone-Siltstone Facies A short walk west of the road brings us to an excellent exposure of the sandstone-siltstone facies (in the rocks of the Grand Wash trough), which exceeds 100 m in thickness. The sandstone-siltstone facies is characterized by alternating beds of pale reddish-brown, moderately sorted, ﬁne- to mediumgrained sandstone, siltstone, and mudstone, with subordinate lenses of pebble conglomerate and gypsum. This unit interﬁngers with both the overlying Hualapai Limestone and underlying fanglomerate facies. The depositional environment of the sandstone siltstone facies was probably a highly evaporative interior continental playa, as evidenced by the intercalated gypsum, thin bedding, and mudcracks (Wallace et al., 2005). The lack of ﬂuvial textures, such as cross-beds or ripple marks, local abundance of gypsum rinds, interﬁngering and bordering fanglomerate facies,

Directions to Stop 1-8 Continue north on Grapevine Wash road and keep left past corral at 0.3 miles. Continue north another 1.3 miles past corral to Stop 1-8.

Brief stop in wash to view fanning west-dipping beds in Hualapai Limestone. Dips decrease appreciably up-section in the limestone, and several beds onlap and pinch out against more steeply dipping beds (Fig. 15). These relations indicate that some extension coincided with deposition of the 11–6 Ma Hualapai Limestone. Directions to Stop 1-9 Continue north on Grapevine Wash Road 3.0 miles to Stop 1-9. Road will descend into the 200–300 m deep Grapevine Canyon and through much of the 300-m-thick Hualapai Limestone. Stop 1-9. Hualapai Limestone and 10.94 Ma Tephra At this stop, a distinctive light gray tephra is interbedded in the lowermost part of the Hualapai Limestone and is well exposed in a conspicuous enclave and bench on the lower east side of the canyon (Fig. 16) only a few meters east of the road. This tephra geochemically correlates with a 10.94 ± 0.03 tuff derived from the Bruneau-Jarbidge volcanic ﬁeld in southernmost Idaho (M. Perkins, 1998, written commun.) and has also yielded an 11.08 ± 0.27 Ma 40Ar/39Ar date on ﬁne-grained sanidine (Faulds et al., 2001c; Wallace et al., 2005). As such, it provides an excellent older age constraint for the Hualapai Limestone. A tephra within the upper part of the limestone on Grapevine Mesa yielded a poorly deﬁned 40Ar/39Ar age of 7.43 ± 0.22 Ma, which should be con-

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Figure 14. Looking northeast at 4.4 Ma basalt of Sandy Point (Tbsp), which is intercalated in Colorado River sediments (QTc and QTcgc).

Figure 15. View north of growth-fault sequence in Hualapai Limestone in Grapevine Wash. West tilts decrease up section from ~25° to subhorizontal in this area. Note onlap of gently dipping beds against more steeply tilted layers. This deformation has been attributed to minor normal faulting (Wallace et al., 2005).

Figure 16. View north of ca. 11 Ma tephra near base of the Hualapai Limestone in Grapevine Canyon. Black arrows point to ~1–2 m light gray tephra beneath ledge of limestone.

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sidered a maximum eruptive age (M. Heizler, 1999, written commun.). A tephra interbedded within the upper part of the Hualapai Limestone in the Temple Bar area ~25 km to the west of the Grand Wash trough yielded an 40Ar/39Ar age of 6.0 Ma (Spencer et al., 2001), which has been widely used as a younger age constraint on the Limestone in the Lake Mead area. The Hualapai Limestone in the Grand Wash trough may therefore be as young as ca. 6 Ma. In both the Gregg Basin and Grapevine Wash area of the Grand Wash trough (Figs. 2 and 5), the Hualapai Limestone appears to form large lenticular wedge-shaped deposits that thicken eastward toward the deeper parts of the half grabens and onlap older units along the western margins of the basins. In the Grapevine Wash area, the limestone consists of a thick (~300 m) section of wavy laminated, vuggy pelleted packstone, wackestone, and calcareous mudstone, with rare siltstone laminations (Wallace, 1999). Petrographic analysis indicates that the limestone is dominantly pelmicrite and rarely contains fossils (Wallace, 1999). In contrast to the interpretation of Blair and Armstrong (1979), detailed analysis of the limestone revealed no evidence of a marine or estuarine setting (Wallace, 1999; Faulds et al., 2001c). Fossils include only ostracodes, green algae (including charophytes), algal mats, oncolites, and stromatolites in a typically pelleted micrite substrate. None of the fossils are diagnostically marine (e.g., Heckel, 1972). Ostracodes are highly tolerant organisms that can live in extreme conditions that include fresh to hypersaline, clear to sediment-loaded water. Blue-green and green algae can also live in several environments but are constrained to the photic zone. However, charophytes need clear, fresh water to survive (Heckel, 1972). Furthermore, locally abundant ﬂowstone indicates a constant source of fresh water super-saturated in calcium carbonate, either issuing from springs or perennial streams. In addition, low δ18O values and highly variable δ13C characterize the Hualapai Limestone (Wallace, 1999), both of which are indicative of a nonmarine setting (e.g., Talbot and Kelts, 1990). The composition and isotopic characteristics of the Hualapai Limestone indicate deposition in one or more restricted warm, shallow, and quiet lakes fed by a relatively continuous source of fresh water in an evaporative climate (Wallace, 1999; Faulds et al., 2001c). In order to stay fresh, the lake or lakes probably had an outlet. DAY 2. KINGMAN TO NORTHERN WHITE HILLS Introduction On Day 2, we will focus on the South Virgin–White Hills detachment fault, synextensional sedimentary deposits in easttilted half grabens (northern White Hills basin) in the hanging wall of the detachment, and upper Miocene paleovalleys connecting the Grand Wash trough–Gregg Basin with the northern White Hills basin and Temple Bar–Detrital basins farther west. Although the Grand Wash fault clearly marks the structural and physiographic boundary between the Colorado Plateau and Basin and Range province in this region, the South Virgin–White Hills detachment fault is considered the most signiﬁcant normal fault in

the eastern part of the Lake Mead domain, as it accommodated as much as 17 km of normal displacement along the west ﬂank of the Gold Butte block directly north of Lake Mead (Brady et al., 2000), about three times the maximum displacement on the Grand Wash fault, and it accommodated >5 km of displacement in the southern White Hills. Thus, the South Virgin–White Hills detachment fault accommodated large-magnitude extension within <30 km of the western margin of the Colorado Plateau and clearly records a progressive westward increase in strain across this region. Directions to Stop 2-1 From the hotel in Kingman, drive north on Andy Devine Boulevard 0.4 miles north to Interstate 40 (I-40). Go west on I-40 1.0 miles and take the Stockton Hill exit. Travel north on Stockton Hill Road 49.5 miles to its end at the intersection with the Pearce Ferry Road. Turn right onto the Pearce Ferry Road and travel 0.4 miles to the east, then turn left (north) onto the Gregg’s Hideout Road. After 3.1 miles, the Gregg’s Hideout Road jogs to the west, so veer to the left, travel 0.9 miles west, then veer back to the right, heading north again on the Gregg’s Hideout Road. Then, continue north 7.3 miles on the Gregg’s Hideout Road, before turning left toward Temple Bar. Go west on the Temple Bar road 6.8 miles, then turn right (north) into Salt Spring Wash. Travel north in Salt Springs Wash 1.7 miles to Stop 2-1 (Fig. 10). Stop 2-1. Salt Spring Detachment, Part of the South Virgin–White Hills Detachment Fault Walk east a few hundred meters up small wash. Note chlorite cataclasite (or breccia) with subhorizontal fabric on wash walls. This is part of the damage zone associated with the detachment. Note its thickness (as you walk up wash) and the gently westdipping mesoscopic faults or shear bands. Kinematic indicators, including asymmetric boudins, sigmoidal foliations, and shear bands within the cataclasite record top-west shear sense, consistent with the east-northeast dip of upper-plate strata. In addition, highly altered (to orange color) and offset basalt dikes corroborate the topwest shear sense. At the top of the wash, a white tuff rests directly on the fault surface. We can debate whether the tuff was deposited on the fault surface or is downfaulted against it. In view of the fact that a 15.2 Ma east-tilted tuff is present in the hanging wall just west of Salt Spring Wash, we prefer the latter interpretation. Within this vicinity, small exposures of a brown microcrystalline rock mark the detachment fault proper. This rock is an ultracataclasite, and we interpret it as analogous to the well-known and thicker “microbreccia ledge” of higher-displacement detachments. Structurally below the detachment surface proper, a zone of greenschist-grade retrogression is present that ranges in thickness from 50 to more than 150 m. This retrogression is superimposed on regional, granulite-facies Proterozoic metamorphism (e.g., Volborth, 1962; Duebendorfer et al., 2001). Although variable, foliation in the footwall of the South Virgin–White Hills detachment generally dips less than 30° to the west in contrast to the generally steep to subvertical dips that characterize crystalline

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary rocks throughout northwestern Arizona (Blacet, 1975; Duebendorfer et al., 2001). The coincidence of gently dipping foliation with pervasive (Miocene?) retrogression of rocks in the lower plate suggests that tilting of the footwall accompanied motion along the South Virgin–White Hills detachment fault. Directions to Stop 2-2 Travel back to the south 0.1 miles in Salt Springs Wash. Stop 2-2. Middle Miocene Synextensional Deposits Take a short traverse through east-dipping conglomerate, sedimentary breccia (all crystalline clasts), and well-bedded sandstone toward a prominent outcrop of white pegmatite. The conglomerates and sedimentary breccias are interpreted as distal debris ﬂow deposits because of their typically nonbedded, matrix-supported character. At the top of a small hill, an outcrop of highly and pervasively shattered pegmatite marks the base of a massive megabreccia sheet that ranges in thickness from 50 to more than 200 m and can be followed along strike for more than 1000 m. This sheet is overlain by an ash-ﬂow tuff dated at 15.2 Ma (40Ar/ 39Ar, sanidine, Duebendorfer and Sharp, 1998; Fig. 17). The megabreccia deposits within this lower sequence of the hanging-wall section are clearly interbedded with stratiﬁed sedimentary rocks and are thus not fractured crystalline rocks of the lower plate that have been somehow infaulted with hanging-wall rocks. The megabreccias exhibit both crackle and jigsaw breccia textures (Yarnold and Lombard, 1989) and are interpreted as catastrophic rock avalanche deposits which are commonly associated with topographic relief generated by active faults (e.g., Yarnold and Lombard, 1989; Topping, 1993), in this case, the South Virgin–White Hills detachment fault.

Th

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Directions to Stop 2-3 Travel back in Salt Springs Wash to the north 0.9 miles. Stop 2-3. Spectacular Exposure of Huge Megabreccia Deposit Fault-like rocks in a megabreccia deposit are exposed on the ridge west of the wash. An underlying debris ﬂow deposit is evident at the level of the wash. Directions to Stop 2-4 Travel south back uphill in Salt Springs Wash 1.5 miles, past road from Greggs Hideout, and continue another 3.0 miles. Stop 2-4. Fault Cataclasite on the Salt Springs Detachment An iron-stained ledge of cataclasite dips gently north here over a footwall of retrograded gneiss in the footwall of the South Virgin–White Hills detachment fault. The footwall here in the Golden Rule Peak area is part of a prominent westward salient, and coincides with a steep gravity gradient over the footwall rocks. The poorly exposed hanging-wall rocks in this area include granite and gneiss as well as unconformably overlying Miocene sedimentary deposits of megabreccia and fanglomerate. Directions to Stop 2-5 Return northward 3.0 miles, passing en route a dipping white ash-fall tuff that is interbedded with fanglomerate and is correlated by Andrei Sarna-Wojcicki with a 12.0 Ma tephra. Turn back to the right on the road toward Greggs Hideout. Proceed 6 miles, partly along the South Virgin–White Hills detachment fault boundary between the dipping section of fanglomerate and 12 Ma tuff juxtaposed against Proterozoic gneiss in the

Gold Butte Block

Tb1

Tc1

Tdt

Figure 17. View to north of Salt Spring Wash. The Gold Butte block, which lies in the lower plate of the South VirginWhite Hills detachment fault, is on the horizon (right distance). White unit in the foreground is a poorly welded ash-ﬂow tuff (Tdt) dated at 15.2 Ma that is interbedded with megabreccia and debris-ﬂow deposits (Tc1) in the upper plate of the South Virgin–White Hills detachment. This tuff probably correlates with the tuff of Mt. Davis (e.g., Faulds et al., 2002b). These rocks are tilted 50–60° east and are unconformably overlain by a 20° northeast-dipping, 14.6 Ma basalt ﬂow (Tb1, dark cap rock in the middle distance). White unit at upper left is the ﬂat-lying, 11–6 Ma Hualapai Limestone (Th).

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footwall of the Salt Spring detachment, and then through the footwall gneisses Stop 2-5. Deposits of Hualapai Wash Walk southwest up short gully to examine a sequence of sandstone and siltstone that underlies cliff-forming cemented fanglomerate (Fig. 18). This sequence may record the initial entry of the Colorado River onto a basin ﬂoored by Hualapai Limestone before the river began incising into the limestone (Howard and Bohannon, 2001). For many miles the sequence can be traced concordantly overlying the gently east-dipping Hualapai Limestone and is tilted gently east with the limestone toward the Wheeler Ridge fault system. The ﬁne-grained deposits in places contain rare rounded pebbles of chert, limestone, and quartzite, clasts typical of Colorado River deposits. Note clayballs and other structures in the sandstone. No other Colorado River ﬂuvial deposits in the Basin and Range province are as high in elevation as these remnants, >700 m. Directions to Stop 2-6 Continue eastward 0.8 miles to Greggs Hideout Road and turn left. Proceed north 7.2 miles to Greggs Hideout on Lake Mead, passing outcrops of Hualapai Limestone en route that are unconformable on the gneisses in the footwall block of the South Virgin–White Hills detachment fault. Stop 2-6. South Virgin–White Hills Detachment Fault at Greggs Hideout The South Virgin–White Hills detachment fault is well exposed at Greggs Hideout as a gouge zone dipping 37° under fanglomerate in the hanging wall (Fig. 19). The fanglomerate in the hanging wall is near-horizontal except within 100 m of the fault, where it bends to 38° dip into the fault and is broken by small-displacement faults. Breccia of Proterozoic rock in the footwall may be either tectonic breccia along the fault zone, or a sliver of megabreccia deposit. Younger horizontal fanglomerate that forms bluffs high to the east overlaps the fault and its footwall. The large boulders in the upper fanglomerate consist of granite closely resembling the Gold Butte Granite. Two kilometers north of here the fault dips more gently over gneiss and granite in the footwall. Directions to Stop 2-7 Return back eastward on the Greggs Hideout Road 1.4 miles to the top of a grade. Stop 2-7. Colorado River Paleochannel and Views of Hualapai Limestone Paleovalley Fill on the Footwall Block Walk northward across lags of rounded river gravel and up a hill underlain by undeformed upper Miocene fanglomerate for views. This site is just south of where the Colorado River crosses the footwall block in Virgin Canyon, and the high-level rounded gravels encountered record an abandoned Pliocene paleovalley of

the river. The northern White Hills basin to the west and southwest is capped by the upper Miocene Hualapai Limestone. The Hualapai Limestone and underlying fanglomerate, which here bridge across the footwall block and connect the White Hills and Greggs basins, postdate the latest motion on the South Virgin– White Hills detachment fault. From this vantage, a paleovalley ﬁll of the Hualapai Limestone is easy to view as inset into the gneisses of the footwall (Fig. 20). The Hualapai Limestone to the east dips gently eastward in the rollover fold against the Wheeler Ridge fault. Spring Canyon, north of the limestone paleovalley ﬁll, contains rubble derived from the limestone that suggests the canyon was also once occupied by the limestone. Rounded river gravels overlying that rubble indicate that the paleovalley was then occupied by a Pliocene course of the Colorado River. The paleovalley and its remnant strings of rubble and river gravels now slopes gently eastward, which suggest that it was back tilted and its drainage direction reversed during development of the rollover fold toward the Wheeler Ridge fault system. DAY 3. KINGMAN TO SOUTHERN WHITE HILLS (THEN RETURN TO LAS VEGAS) Introduction Day 3 will involve an overview of the Hualapai basin and an east to west traverse across the southern White Hills basin. The Hualapai basin contains a 2.5-km-thick, middle to late Miocene halite deposit, one of the thickest known nonmarine salt deposits (Faulds et al., 1997). However, because the Hualapai basin is a closed depression that has not been dissected by tributaries to the Colorado River, synextensional deposits are not exposed. Interpretations for this basin are therefore based on geophysical and well data. Two stops will provide an overview of the Hualapai basin and neighboring northern Cerbat Mountains. We will then cross the west-dipping Cerbat Mountains–Cyclopic fault (southern leg of the South Virgin–White Hills detachment fault) and enter into the southern White Hills basin, where several stops will facilitate discussion of the timing of extension and paleogeographic evolution of the region. The conglomerate of Rock Spring dominates the eastern subbasin of the southern White Hills and records signiﬁcant erosion of the footwall of the South Virgin–White Hills detachment fault and possibly the western margin of the Colorado Plateau. It will therefore be the focus of several stops. En Route Discussion The eastern part of Kingman resides near the southwestern margin of the Hualapai basin. As we leave Kingman and travel north on the Stockton Hill Road, the Hualapai basin will open up to the east (on our right) and the Cerbat Mountains will dominate to the west. The southern Grand Wash fault bounds the Hualapai basin on the east. The Grand Wash Cliffs and Colorado Plateau lie in the footwall of the southern Grand Wash fault. The western margin of the Colorado Plateau is more highly dissected east of

18. Stop 2-5. Cliff-forming 143 ceCenozoic evolution of the abrupt Colorado Plateau–Basin and RangeFigure boundary

mented fanglomerate consistently overlies the deposits of Hualapai Wash, which consist of sandstone, siltstone, claystone, and rare rounded pebbles of limestone, chert, and quartzite. The deposits of Hualapai Wash may record the initial entry of the Colorado River into the Basin and Range province. The deposits of Hualapai Wash and the overlying cemented fanglomerate can be traced for many kilometers northward toward Gregg Basin, everywhere concordantly overlying the folded Hualapai Limestone. The deposits of Hualapai Wash therefore were deposited on the uppermost Miocene basin-ﬁlling sequence. Younger Colorado River deposits are inset into Hualapai Limestone and record incision resulting from exterior drainage and lowered base levels.

Figure 19. Detachment fault at Greggs Hideout (Stop 2-6), marked by black arrow. Foliated gouge separates a hanging wall of internally faulted Miocene fanglomerate from a footwall of breccia, consisting primarily of Proterozoic rock.

Grand Wash Cliffs

Th

Th

Figure 20. View eastward from Stop 2-7 at light-toned Hualapai Limestone (Th) ﬁlling paleotopography cut into dark Proterozoic gneiss in the northern White Hills. The limestone and local underlying upper Miocene conglomerate (red slopes in foreground) here bridge across a low part of the block that forms the footwall of the South Virgin–White Hills detachment fault. The sediments thus record a time when these two basins ﬁlled to the point of merging, and drainage basins became more integrated. A younger paleovalley of the Pliocene Colorado River followed a nearby path.

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the Hualapai basin than it is to the north directly east of the Grand Wash trough. This morphologic difference probably results from more widespread exposure of less resistant Proterozoic rocks to the east of the Hualapai basin in contrast to the highly resistant Cambrian through Mississippian carbonates that lie directly east of the Grand Wash trough. The aforementioned gentle regional northeast tilt of strata, inherited from early Tertiary time, accounts for the southward trend toward older exposed strata. To our west, the Cerbat Mountains are primarily composed of Paleoproterozoic gneiss and granite (Duebendorfer et al., 2001) but do contain scattered exposures of the basal part of the Tertiary section, especially on the north, east, and southern ﬂanks, as discussed at Stop 3-1. Several gently east-dipping cuestas, capped by resistant volcanic units near or at the base of the Miocene section, grace the east slopes of the Cerbat Mountains and will be quite conspicuous along our route. Directions to Stop 3-1 From the hotel in Kingman, drive north on Andy Devine Boulevard 0.4 miles to I-40. Go west 1.0 miles on I-40 and take the Stockton Hill exit. Travel north on Stockton Hill Road 34.3 miles to Stop 3-1. Stop 3-1. Overview of Northern Cerbat Mountains The Cerbat Mountains are the upthrown western part of the east-tilted fault block that forms the Hualapai basin. Miocene strata within the Cerbat Mountains rest nonconformably on Paleoproterozoic gneiss. The base of the Miocene section is well exposed in the high rugged ridges to our west and consists, in ascending order, of an ~10 m thick basal arkosic conglomerate, thin sequence of ca. 20 Ma basalt ﬂows, and ~250 m of matrixsupported, volcaniclastic conglomerate and intercalated trachybasaltic andesite to trachydacite lavas. A ﬂow near the top of this sequence yielded an 40Ar/ 39Ar age of 18.3 Ma. The volcaniclastic conglomerate comprises the bulk of the high jagged ridges, as well as the prominent pinnacles that dominate the skyline east of the town of Dolan Springs. These ridges contain some of the best exposures of the base of the Miocene section in the northern Colorado River extensional corridor. The 18.5 Ma Peach Springs Tuff, which is well-exposed near Kingman, pinches out northward across the Cerbat Mountains and has not been observed in the northern part of the range. All strata in the northern Cerbat Mountains dip gently east (~15°–22°) and are cut by a series of moderately to steeply west- to southwest-dipping normal faults. Directions to Stop 3-2 Continue north on the Stockton Hill Road 7.6 miles. We will climb to the top of small ridge directly west of the road for an overview of the Hualapai basin. Stop 3-2. Overview of Hualapai Basin and Red Lake Playa The Hualapai basin contains one of the thickest known, nonmarine halite deposits in a continental rift, a 2.5-km-thick section

of middle to late Miocene salt (Faulds et al., 1997). As evident by Red Lake Playa directly east of the Stockton Hill Road, the Hualapai basin is a closed, internally drained depression. The Gregg Basin to the left drains northward to the Colorado River. A very low saddle separates the Gregg and Hualapai basins. Integration of the Hualapai basin into the Colorado River drainage is probably imminent. The salt deposit within the Hualapai basin is not exposed. Saline groundwater noted by ranchers provided the ﬁrst evidence of subsurface salt. Oil companies drilled the basin in the 1950s in hopes of ﬁnding hydrocarbons trapped by a presumed subsurface salt dome. Subsequently, the salt was studied as a potential natural gas repository by the El Paso Natural Gas Company (F.E.R.C., 1982), who drilled additional deep holes in the basin and acquired seismic reﬂection proﬁles. Seismic reﬂection and drill-hole data indicate that the littledeformed, unexposed halite is 2.5-km-thick in the central part of the basin, approaches ~200 km3 in volume, and has a three-dimensional lenticular-wedge geometry (Fig. 13). An age of 13–8 Ma is suggested for the salt because it lies in the upper, more gently tilted part of a growth-fault sequence, and extension in the region is bracketed between ca. 16 and 8 Ma. The texture and bromine content of the halite, dominance of halite, and S and O isotopic values of intercalated and capping anhydrite indicate that halite deposition took place in an intracontinental playa that accommodated regional groundwater discharge. Several events conspired to produce this unusually thick salt deposit, including regional aridity, a broad catchment basin with a closed drainage network, ample supplies of Na+ and Cl–, and rapidly developing accommodation space within the basin (Faulds et al., 1997). Large evaporite bodies of probable nonmarine origin are common in many of the basins of southern and central Arizona (Peirce, 1976). The ridges to the west of the road consist of Paleoproterozoic gneiss capped by ca. 18 Ma basaltic trachyandesite ﬂows. These ca. 18 Ma basaltic andesite lavas may correlate with a southwestward thickening sequence of lower Miocene lavas of similar age and composition in the Iron Mountain and Garnet Mountain areas (e.g., 17.4 Ma—K/Ar age reported in Lucchitta and Young 1986; Wenrich et al., 1996) along the western margin of the Colorado Plateau. Directions to Stop 3-3 Continue north 7.6 miles on the Stockton Hill Road, then turn left onto the Pearce Ferry Road. Travel 4.6 miles on the Pearce Ferry Road and pull off to the left into a gravelly parking area at the foot of a small hill. We will take a short traverse to the top of this hill. Stop 3-3. Cerbat Mountains Fault A short walk will take us to the trace of the Cerbat Mountains fault, which is marked by the contact between 8.7 Ma olivine basalt on the west and Paleoproterozoic gneiss on the east. The Cerbat Mountains fault bounds the southern White Hills basin on the southeast and strikes north-northeast through the saddle

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary between the southern White Hills and northern Cerbat Mountains. In the footwall of the fault to our east, ca. 18 Ma maﬁc lavas at the base of the Miocene section cap Paleoproterozoic gneiss, whereas in the hanging to the west, the 8.7 Ma basalts of the Table Mountain Plateau cap a 3-km-thick Miocene section in the southern White Hills basin. The offset of the early Miocene nonconformity between the Table Mountain Plateau area and northern Cerbat Mountains suggests ~4–6 km of normal separation along the northern part of the Cerbat Mountains fault (Fig. 11B). The Cerbat Mountains fault probably dips moderately west in its upper part but has an overall listric geometry, as evidenced by signiﬁcantly greater tilting of the hanging wall, as compared to the footwall (e.g., as much as 50° in the Table Mountain Plateau area versus ~20° in the northern Cerbat Mountains). A Proterozoic gneiss in the footwall of the Cerbats Mountains fault yielded an apatite ﬁssion-track age of ca. 73 Ma, whereas all crystalline rocks in the footwall of the South Virgin–White Hills detachment fault to the north yield apatite ﬁssion-track ages of <20 Ma. The Cerbat Mountain fault merges with the Cyclopic fault (e.g., Blacet, 1975; Myers et al., 1986) directly west of Table Mountain Plateau (Figs. 3 and 11A). Collectively, the two faults deﬁne a three-dimensional, scoop-shaped geometry of the southern White Hills basin (Fig. 11). There is no evidence that one fault accommodated offset of the other. It is therefore likely that the Cerbat Mountains fault represents a southern continuation of both the Cyclopic fault and South Virgin–White Hills detachment fault. The overall length of this fault system from the Gold Butte block on the north to the southern Cerbat Mountains on the south is at least 140 km (Fig. 2). A northward increase in both the breadth and metamorphic grade of the footwall crystalline terrane and a corresponding change in fault rocks from cataclasites to mylonites indicate a signiﬁcant northward increase in displacement on the South Virgin–White Hills detachment fault between the southern White Hills and Gold Butte block (Duebendorfer and Sharp, 1998). In the South Virgin Mountains, the Lakeside Mine segment of the South Virgin–White Hills detachment may have accommodated as much as 17 km of normal displacement (Brady et al., 2000). North of the southern White Hills, this fault zone accommodates the greatest amount of normal displacement within the Lake Mead region. To the south of the southern White Hills, this fault zone essentially borders the eastern edge of the highly extended extensional corridor, as it separates the gently tilted Cerbat Mountain block from highly extended terrane in the Black Mountains. This fault zone is clearly one of the most signiﬁcant structures in the northern Colorado River extensional corridor. Directions to Stop 3-4 Cross Pearce Ferry Road obliquely and turn right onto dirt road, marked by a stop sign at the intersection with the highway (<0.1 miles from Stop 3-3). Travel west on dirt road and veer right around corral after 0.6 miles. Continue 1.6 miles on dirt road to Stop 3-4.

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Stop 3-4. View of Faulted 8.7 Ma Basalts This stop provides a good view to the south of the gently (~5°–10°) east-tilted 8.7 Ma basalts that cap the Table Mountain Plateau, which are as much as ~250 m thick. The basalts are cut by several west-dipping normal faults and the Cerbat Mountains fault, indicating that minor extension continued to at least ca. 8 Ma. It is possible that the faults cutting the basalt were active contemporaneously with the Wheeler Ridge fault and Lost Basin Range faults to the east (i.e., since 4.4 Ma). The 8.7 Ma basalts overlie the conglomerate of Rock Spring, with little if any angular discordance between the conglomerate and basalts. Other documented lavas of this age in the region cap Callville Mesa in the western Lake Mead area (Feuerbach et al., 1993) and Snap Point along the northern Grand Wash Cliffs on the westernmost margin of the Colorado Plateau (Fig. 2; Faulds et al., 2001b). The basalts of Table Mountain Plateau probably ﬁlled paleochannels cut into the underlying conglomerate of Rock Spring and appear to have pooled in the lowermost part of the southern White Hills basin. They were subsequently tilted eastward, probably in response to movement along the Cerbat Mountains fault. Owing to the highly resistant basalts, the present topography at Table Mountain Plateau (Fig. 12) is clearly inverted from that in late Miocene time. Directions to Stop 3-5 As we make our way to the west in the southern White Hills basin between Stops 3-5 and 3-6, we will traverse down-section through the synextensional ca. 16.7–8.7 Ma conglomerate of Rock Spring, with numerous small road cuts exposing the basement-clast conglomerate. It is important to note that this area is currently under development; the network of roads may change signiﬁcantly with further development. From Stop 3-5, continue on dirt road 0.6 miles to the north and turn left (west) at major cross road. After 1.0 miles, go straight at oblique cross road, then continue for another 2.3 miles and keep right at Y-intersection. After another 0.9 miles, keep right again after another Y-intersection, then continue 0.6 miles to Stop 3-6, located in a small pass at the crest of the White Hills, where recent excavations have left a large parking area. Prepare for a short hike. Stop 3-5. Conglomerate of Rock Spring The recent excavations at the pass have resulted in a superbly exposed knob of the Rock Spring conglomerate, which still existed during planning of this ﬁeld trip in September 2007. The matrix supported texture, large clasts of megacrystic Proterozoic granite, and moderate east dips are readily visible in the exposure (Fig. 21). Intercalated in the conglomerate are lenses of basaltic andesite lava and tuff, including the 15.2 Ma tuff of Bridge Spring (cf. Faulds et al., 2002b). Clasts of Proterozoic gneiss and granite dominate the conglomerate of Rock Spring and were largely derived from the footwall of the South Virgin–White Hills

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Figure 21. Conglomerate of Rock Spring. View looking south at poorly sorted, matrix-supported conglomerate of Rock Spring, containing abundant clasts of Paleoproterozoic megacrystic granite. Large boulder of megacrystic granite at top of mound is ~5 m long. White dashed lines denote faint bedding, which dips moderately east. Black arrow points to relatively distinct bedding plane. Small white oval surrounds 40-cm-long rock hammer.

detachment fault (Cyclopic and Cerbat Mountains segments). Abundant clasts of Proterozoic megacrystic granite throughout the ca. 16.7–8.7 Ma section of conglomerate further suggest that the Garnet Mountain area along the western margin of the Colorado Plateau may have been a source for some of the detritus. These relations indicate signiﬁcant headward erosion into the footwall of the South Virgin–White Hills detachment fault and possibly the western margin of the Colorado Plateau by westﬂowing drainages as early as middle Miocene time. A short ~0.5 mile walk to the south will bring us to a nice vantage point, with panoramic views of the southern White Hills. Appreciable tilt fanning can be observed to the south, whereby the 15.2 Ma tuff of Bridge Spring is tilted ~35° east in a prominent hogback and the capping 8.7 Ma basalts at Table Mountain Plateau dip ~5° east. Dips in the conglomerate in this area, slightly down section of the tuff of Bridge Spring, are ~55°. Tilt fanning within the southern White Hills indicates that major east-west extension occurred from ca. 16.7–8 Ma. As we traverse through exposures of the Rock Spring conglomerate, note the abundant clasts of garnet-bearing gneiss and megacrystic granite, as well as a thin basaltic andesite ﬂow. To the north, the prominent peak of Senator Mountain can be seen. Senator Mountain is capped by ca. 9.9 Ma basalt ﬂows that slope gently west but overlie the gently east-tilted Rock Spring conglomerate. A north-striking basalt dike directly east of Senator Mountain probably fed these ﬂows, which ﬂowed westward in a paleochannel, which has now been inverted due to the resistant basalts. Directions to Stop 3-6 From Stop 3-5, continue west on dirt road. Note Buick-size clast of megacrystic granite directly north of road after 0.5 miles. Continue another 0.6 miles to west to old water tank and Stop 3-6.

Stop 3-6. Lower Part of Rock Spring Conglomerate On the south side of the wash, the conglomerate of Rock Spring interﬁngers with a thin sequence of basaltic andesite lavas, which have yielded an 40Ar/ 39Ar age of ca. 16.0 Ma. The volcanic component is clearly subordinate to sedimentary rocks in the eastern subbasin of the southern White Hills basin. Exposures here lie in the lower part of the Rock Spring conglomerate, only a few hundred meters east of the nonconformity with Proterozoic basement. Directions to Stop 3-7 Continue west on dirt road and keep left (southwest) at Yintersection 0.4 miles from Stop 3-6. After another 0.2 miles, road veers west and soon passes into Proterozoic gneisses, which directly underlie the Rock Spring conglomerate in this area. Continue west another 1.1 miles, then keep left at Y-intersection. The road will then cross the approximate trace of the Mountain Spring fault, which is a major west-dipping normal fault that divides the southern White Hills into two discrete subbasins (western volcanic-dominated and eastern sedimentary-dominated subbasins). A narrow northerly trending late Tertiary basin lies in the hanging wall of the Mountain Spring fault in this area, within which the small community of White Hills resides. Displacement on the fault in this area may exceed ~4 km, but decreases appreciably to the south. The Mountain Spring fault can be considered a splay of the Cyclopic–Cerbat Mountains fault system, as it links northward with the Cyclopic fault near Senator Mountain. It may therefore help to accommodate the southward decrease in slip on the South Virgin–White Hills detachment fault. About 0.9 miles past the Y-intersection, turn right and then immediately left onto wide dirt road. After 0.4 miles, wide road curves 90° to the west into Indian Peak Drive. After traveling

Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary 1.0 miles west, turn left onto Senator Boulevard and continue south for 1.2 miles, then veer left onto Skipper Boulevard. Travel 0.4 miles on Skipper Boulevard, then turn right (west) onto paved road of White Hills Boulevard. White Hills Boulevard descends through a small canyon. The surrounding ridges are composed primarily of basaltic andesite lavas tilted ~25°–35° to the east, which have yielded ca. 16 Ma 40Ar/39Ar ages. To the north of the road, gently (~10°) east-tilted Late Tertiary conglomerate and sandstone onlaps the maﬁc lavas on the east side of the ridge. From the intersection with Skipper Boulevard, travel west 1.0 miles on White Hills Boulevard to Stop 3-7, where we will take a very short hike to the north of the road. Stop 3-7. Early Miocene Nonconformity and Volcanic Rocks The steeply east-tilted, early Miocene nonconformity (i.e., base of Miocene section) is exposed ~100 m north of the road on the lower west ﬂank of the high north-trending ridge. A dacite lava near the base of the section is ca. 18.5 Ma. The dacite-andesite lavas correlate with the volcanics of Dixie Queen Mine, a thick (2–3 km) section of intermediate lavas found in the Mount Perkins area to the west of the Detrital basin (Faulds et al., 1995). Younger, more gently tilted (~30°), ca. 16 Ma basaltic andesite lavas onlap the steeply tilted (~75°) volcanics of Dixie Queen Mine on the ridge to the east. The capping maﬁc lavas to the east are part of a thick synextensional, bimodal volcanic sequence that dominates the western subbasin of the southern White Hills in contrast to the sediment-dominated eastern subbasin. Beneath the nonconformity to the northwest, extensive mine workings in the Proterozoic gneiss are the remnants of a mining boom in the White Hills more than 100 years ago. The boom town of White Hills was located in this area. The Hualapai and Paiute Indians had long used the iron and manganese oxides associated with veins in the White Hills for paints. This was the Indian Secret mining district, or Silverado district, where gold and silver (“horn silver” or “chloride silver”) was produced from 1892 to 1899 (Huskinson, 1984). Total production was probably 6–8 million ounces of silver and as much as 5000 ounces of gold. In 1899, however, the town of White Hills was destroyed in a ﬂash ﬂood with signiﬁcant loss of life. The ﬂood still marks one of the worst catastrophes in Arizona history. End of Trip, Return to Las Vegas Continue west on White Hills Road 6.1 miles to U.S. Highway 93. Turn right onto Highway 93 and proceed north to Hoover Dam and Las Vegas. SUMMARY An unusually abrupt boundary separates the Colorado Plateau and Basin and Range province in northwestern Arizona. Little deformed, subhorizontal strata along the western margin of the Colorado Plateau give way westward to moderately to steeply east-tilted fault blocks within the northern Colorado

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River extensional corridor across a system of major west-dipping normal faults, which include the Grand Wash and South Virgin–White Hills detachment fault zones. The Grand Wash fault zone forms the main structural and physiographic boundary between the Colorado Plateau and extensional corridor, with as much as 5 km of displacement across some segments. However, the South Virgin–White Hills detachment fault accommodates the greatest amount of extension in the region, having accumulated 5–17 km of normal displacement <30 km to the west of the plateau margin. Several large basins (east-tilted half grabens) developed in the hanging walls of the Grand Wash and South Virgin–White Hills detachment faults. Foundering of the extensional corridor relative to the Colorado Plateau helped to induce excavation of at least the western part of the Grand Canyon, which was carved into the Colorado Plateau directly east of the corridor. Cenozoic strata within the hanging walls of the Grand Wash and South Virgin–White Hills detachment faults thus afford an opportunity to assess both the timing of structural demarcation between the Basin and Range and Colorado Plateau and development of the Colorado River. On this trip, we viewed stratigraphic sections and structures in both the hanging wall of the Grand Wash fault (Grand Wash trough and Gregg and Hualapai basins) and hanging wall of the South Virgin–White Hills detachment fault (northern and southern White Hills basins). Tilt fanning within the basins indicates that major eastwest extension began ca. 16.5 Ma and had eased signiﬁcantly by ca. 13 Ma. A northwestward shift of major extension into the western Lake Mead region 13–9 Ma further lowered local base-level in the Basin and Range. However, minor extension continued to at least the early Pliocene and possibly Pleistocene time in some areas, as best evidenced by 300 m of post–6 Ma down-to-the-west displacement on the Wheeler Ridge fault in the Gregg basin (Fig. 8). These timing constraints indicate that the abrupt boundary between the Colorado Plateau and Basin and Range province in northwest Arizona was taking shape by ca. 16 Ma and had largely developed by ca. 13 Ma. Footwall blocks, including the western margin of the Colorado Plateau, were shedding detritus into the basins as early as ca. 16 Ma. This implies that streams began eroding headward into the Colorado Plateau in the middle Miocene as base level lowered to the west. Excavation of the western Grand Canyon by a small west-ﬂowing stream may have also begun in the middle Miocene. However, thick middle to late Miocene evaporite and lacustrine deposits (e.g., Hualapai Limestone) demonstrate that a long period of internal drainage followed the onset of major extension. The widespread distribution of such deposits and their proximity to the present course of the Colorado River (Fig. 9) may signify a large inﬂux of groundwater and/or surface waters from developing drainage systems, both of which may have been related to subsequent development of the Colorado River. Stratigraphic relations indicate, however, that a through going Colorado River did not arrive in the Lake Mead region until ca. 5.6–4.4 Ma.

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ACKNOWLEDGMENTS This work was funded by National Science Foundation grants awarded to Faulds (EAR93-16770, EAR99-10977, and EAR-0409913) and Duebendorfer (EAR-9909275 and EAR0424900), as well as EDMAP cooperative agreements 1434HQ-96-AG-01546 and 1434-HQ-97-AG-07146) provided by the U.S. Geological Survey (both awarded to Faulds). Duebendorfer gratefully acknowledges support from the donors of the Petroleum Research Fund, administered to the American Chemical Society (PRF 29331-B2). Partial ﬁeld support was also provided by the U.S. Geological Survey in Las Vegas, Nevada, for which we thank Gary Dixon and Pete Rowley. In addition, the National Park Service at the Lake Mead National Recreation Area has provided critical logistical support for some of this work; special thanks to Kent Turner and Darlene Carnes for this support. Several graduate students have contributed signiﬁcantly in recent years to our understanding of this region, including Linda Price, Mark Wallace, Nathan Blythe, Brian Coven, Pat Kelleher, and Zack Swaney. In addition, we have beneﬁted greatly from discussions with Sue Beard, Luis Gonzalez, Scott Lundstrom, Dwight Schmidt, Eugene Smith, and Paul Umhoefer. We also thank Terri Garside for help with setting up logistics for the trip and Kris Pizarro for assistance in preparing some of the ﬁgures. Constructive reviews by Jon Spencer and Gene Smith improved this manuscript. REFERENCES CITED Anderson, R.E., 1971, Thin skin distention in Tertiary rocks of southeastern Nevada: Geological Society of America Bulletin, v. 82, p. 43–58, doi: 10.1130/0016-7606(1971)82[43:TSDITR]2.0.CO;2. Anderson, R.E., 1978, Geologic map of the Black Canyon 15-minute Quadrangle, Mohave county, Arizona and Clark County, Nevada, U.S. Geological Survey Geologic Quadrangle Map GQ-1394, scale 1:62,500. Anderson, R.E., and Barnhard, T.P., 1993, Aspects of three-dimensional strain at the margin of the extensional orogen, Virgin River depression area, Nevada, Utah, and Arizona: Geological Society of America Bulletin, v. 105, no. 8, p. 1019– 1052, doi: 10.1130/0016-7606(1993)105<1019:AOTDSA>2.3.CO;2. Anderson, R.E., Longwell, C.R., Armstrong, R.L., and Marvin, R.F., 1972, Signiﬁcance of K-Ar ages of Tertiary rocks from the Lake Mead region, Nevada-Arizona: Geological Society of America Bulletin, v. 83, p. 273– 288, doi: 10.1130/0016-7606(1972)83[273:SOKAOT]2.0.CO;2. Anderson, R.E., Barnhard, T.P., and Snee, L.W., 1994, Roles of plutonism, midcrustal ﬂow, tectonic rafting, and horizontal collapse in shaping the Miocene strain ﬁeld of the Lake Mead area, Nevada and Arizona: Tectonics, v. 13, no. 6, p. 1381–1410, doi: 10.1029/94TC01320. Beard, L.S., 1996, Paleogeography of the Horse Spring Formation in relation to the Lake Mead fault system, Virgin Mountains, Nevada and Arizona, in Beratan, K.K., ed., Reconstructing the history of Basin and Range extension using sedimentology and stratigraphy: Geological Society of America Special Paper 303, p. 27–60. Beard, L.S., Anderson, R.E., Block, D.L., Bohannon, R.G., Brady, R.J., Castor, S.B., Duebendorfer, E.M., Faulds, J.E., Felger, T.J., Howard, K.A., Kuntz, M.A., and Williams, V.S., 2007, Preliminary geologic map of the Lake Mead 30′ × 60′ quadrangle, Clark County, Nevada, and Mohave County, Arizona: U.S. Geological Survey Open-File Report 2007-1010, 109 p., 3 plates, scale 1:100,000 (http://pubs.usgs.gov/of/2007/1010/). Blacet, P.M., 1975, Preliminary geologic map of the Garnet Mountain quadrangle, Mohave County, Arizona: U.S. Geological Survey Open-File Report 75-93. Blackwelder, E., 1934, Origin of the Colorado River: Geological Society of America Bulletin, v. 45, p. 551–566.

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Mills, J.G., 1994, Geologic map of the Hoover Dam Quadrangle: Nevada Bureau of Mines and Geology Map 102, 1:24,000 scale. Morikawa, S.A., 1994, The geology of the tuff of Bridge Spring, southern Nevada and northwestern Arizona [M.S. thesis]: Las Vegas, University of Nevada, 165 p. Myers, I.A., Smith, E.I., and Wyman, R.V., 1986, Control of gold mineralization at the Cyclopic Mine, Gold Basin District, Mojave County, Arizona: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 81, p. 1553–1557. Nations, J., 1989, Cretaceous history of northeastern and east-central Arizona, in Jenney, J.P., and Reynolds, S.J., eds., Geologic evolution of Arizona: Arizona Geological Society Digest, v. 17, p. 435–446. Nielson, J.E., Lux, D.R., Dalrymple, G.B., and Glazner, A.F., 1990, Age of the Peach Springs Tuff, southeastern California and western Arizona: Journal of Geophysical Research, v. 95, p. 571–580. Parsons, T., and McCarthy, J., 1995, The active southwest margin of the Colorado Plateau: Uplift of mantle origin: Geological Society of America Bulletin, v. 107, no. 2, p. 139–147, doi: 10.1130/0016-7606(1995)107<0139: TASMOT>2.3.CO;2. Pederson, J.L., and Karlstrom, K.E., 2001, Relating differential incision of Grand Canyon to slip along the Hurricane/Toroweap fault system, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 159–165. Pederson, J., Karlstrom, K., Sharp, W., and McIntosh, W., 2002, Differential incision of the Grand Canyon related to Quaternary faulting—Constraints from U-series and Ar/Ar dating: Geology, v. 30, no. 8, p. 739–742, doi: 10.1130/0091-7613(2002)030<0739:DIOTGC>2.0.CO;2. Peirce, H.W., 1976, Tectonic signiﬁcance of Basin and Range thick evaporite deposits: Arizona Geological Society Digest, v. 10, p. 325–339. Peirce, H.W., 1985, Arizona’s backbone: the transition zone, Arizona Bureau of Geology and Mineral Technology: Fieldnotes, v. 15, p. 1–6. Powell, J.W., 1875, Exploration of the Colorado River of the West and its tributaries: Explored in 1869, 1870, 1871, and 1872: Washington, D.C., U.S. Government Printing Ofﬁce, 291 p. Powell, J.E., 1895, Canyons of the Colorado River: Meadville, Pennsylvania, Flood and Vincent, The Chautauqua-Century Press, 400 p. Price, L.M., 1997, Geometry and evolution of a major segment of the Grand Wash fault zone, southern White Hills, northwestern Arizona [M.S. thesis]: University of Iowa, Iowa City, 147 p. Price, L.M., and Faulds, J.E., 1999, Structural development of a major segment of the Colorado Plateau-Basin and Range boundary, southern White Hills, Arizona: Nevada Petroleum Society Guidebook, p. 139–170. Reiners, P.W., 2002, (U-Th)/He chronometry experiences a renaissance: Eos (Transactions, American Geophysical Union), v. 83, no. 3, p. 21–27. Reiners, P.W., Brady, R., Farley, K.A., Fryxell, J.E., Wernicke, B., and Lux, D., 2000, Helium and argon thermochronometry of the Gold Butte block, South Virgin Mountains, Nevada: Earth and Planetary Science Letters, v. 178, p. 315–326. Roskowski, J.A., Patchett, P.J., Pearthree, P.A., Spencer, J.E., Faulds, J.E., and Reynolds, A.C., 2007, A late Miocene-early Pliocene chain of lakes fed by the Colorado River: Evidence from Sr isotopes of the Bouse Formation between Grand Canyon and Gulf of California: Geological Society of America Abstracts with Programs, v. 39, no. 6, p. 435. Sherrod, D., and Nielson, J., eds., 1993, Tertiary stratigraphy of highly extended terranes, California, Arizona, and Nevada: U.S. Geological Survey Bulletin 2053. Silver, L.T., Bickford, M.E., Van, S.W.R., Anderson, J.L., Anderson, T.H., and Madaris, L.G., 1977, The 1.4–1.5 b.y. transcontinental anorogenic plutonic perforation of North America: Geological Society of America Abstracts with Programs, v. 9, no. 7, p. 1176–1177. Smith, P.B., 1970, New evidence for a Pliocene marine embayment along the lower Colorado River area, California and Arizona: Geological Society of America Bulletin, v. 81, p. 1411–1420, doi: 10.1130/00167606(1970)81[1411:NEFAPM]2.0.CO;2. Spencer, J.E., and Patchett, P.J., 1997, Sr isotope evidence for a lacustrine origin for the upper Miocene to Pliocene Bouse Formation, lower Colorado River trough, and implications for timing of Colorado Plateau uplift: Geological Society of America Bulletin, v. 109, p. 767–778, doi: 10.1130/0016-7606(1997)109<0767:SIEFAL>2.3.CO;2. Spencer, J.E., and Pearthree, P.A., 2001, Headward erosion versus closed-basin spillover as alternative causes of Neogene capture of the ancestral Colo-

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MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 11 2008

Interpretation of Pleistocene glaciation in the Spring Mountains of Nevada: Pros and cons Jerry Osborn Department of Geoscience, University of Calgary, Calgary, Alberta T2N 1N4, Canada Matthew Lachniet Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154, USA Marvin (Nick) Saines Saines Environmental Hydrogeology, 1587 Figueroa Drive, Las Vegas, Nevada 89123, USA

ABSTRACT There is a long history of debate over glacial versus non-glacial interpretations of both Quaternary and pre-Quaternary diamicts in various places around the world, and the Spring Mountains in southern Nevada are the site of one such debate. Here the debate focuses not only on Quaternary diamicts, but also on landforms and erosional features. The deposits and geomorphic features in question will be examined on this ﬁeld trip. The Spring Mountains are developed in a fault block in the southern Basin and Range Province; elevations range from ~4000 ft (1220 m) at the eastern base to 11,918 ft (3634 m) at Charleston Peak. The range is farther south than any other glaciated range in Nevada; however, the glaciated San Gorgonio Mountains on the border of the Basin and Range in California are farther south, though in a much more maritime position. The Spring Mountains lie in the rain shadow of the Sierra Nevada; rainfall increases from ~4 in (10 cm) per year in the Las Vegas Valley to ~20 in (50 cm) per year at the crest of the range. A published interpretation of sedimentary and geomorphic features at the head of Kyle Canyon claims that steep valley heads of Kyle Canyon and Big Falls wash are degraded cirques, that a ridge at the mouth of Big Falls wash is a lateral moraine, and that diamicts exposed in the ridge include glacial till. An alternative view is that the “cirques” are normal valley heads as are found in high-relief desert ranges, that the “lateral moraine” owes some of its ridge character to erosion along Big Falls wash and may originally have been a debris ﬂow levee or a protalus rampart, and that the “till” is actually colluvium. Abundant clast striations constitute a key element of the glacial interpretation, and much rests on whether glacial striations can be distinguished from mass movement striations. Keywords: Spring Mountains, glaciation, mass movement, striations, Quaternary. Osborn, J., Lachniet, M., and Saines, M., 2008, Interpretation of Pleistocene glaciation in the Spring Mountains of Nevada: Pros and cons, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 153–172, doi: 10.1130/2008.ﬂd011(07). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION This ﬁeld trip focuses on enigmatic alpine landforms and deposits that have been interpreted in at least a couple of ways, one of which is as an indicator of Pleistocene glaciation, and the other of which is as a record of non-glacial surﬁcial processes in high-relief limestone terrain. The landforms and deposits in question are at the head of Kyle Canyon in the Spring Mountains just northwest of Las Vegas (Figs. 1 and 2). Determination of glacial or non-glacial origin of the features will ﬁll a gap in the Pleistocene glacial equilibrium-line altitude (ELA) map of the Great Basin (e.g., Zielinski and McCoy, 1987), which is poorly resolved and complicated by an unusually large Pleistocene/modern ELA difference in the San Bernardino Mountains (Fig. 1) (Owen et al., 2003). The debate in the Spring Mountains hinges mainly on the signiﬁcance of clast striations, and has had the consequence of focusing renewed attention on ways such striations can be generated in non-glacial environments. A complete and/or satisfactory explanation of the deposits in question remains elusive; consequently, the point of this trip is not to present well-established conclusions, but rather to encourage discussion and debate and possibly generate some new ideas.

INTERPRETATION OF GLACIATION: THE GENERAL PROBLEM Glacial erosional and depositional imprints on landscapes are very obvious where glaciers were recently or still are present, but interpretation of past glaciation can be difﬁcult where glaciers were small and/or existed long ago. Most arguments over glacial interpretations have involved pre-Quaternary, particularly Precambrian sediments (general discussion by Schermerhorn, 1974; see also Eyles, 1993); such arguments are crucial to determination of general climatic and/or geochemical states of the planet in ancient times. For example, the Proterozoic Toby Formation in western Canada has been the focus of anti-glacial mappers (e.g., Reesor, 1973) and pro-glacial mappers (e.g., Aalto, 1971), although more recently the glacial hypothesis seems to reign unchallenged. But debates have also focused on late Cenozoic deposits. An example is the “Deadman Pass till” in California, which was claimed by Curry (1966) to be the oldest (3 Ma) Pleistocene glacial deposit known from temperate latitudes, but reinterpreted by Bailey et al. (1990) to be a residual lag and colluvial deposit formed by weathering of poorly consolidated Pliocene pyroclastic rocks that happened to incorporate granitic and metamorphic basement clasts.

Figure 1. Regional physiography and sites mentioned in the text. The dashed outline of Clark County, Nevada, is shown in Figure 2.

Pleistocene glaciation in the Spring Mountains of Nevada

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Figure 2. Shaded relief map of southern Nevada, and location of the Las Vegas Valley, the Spring Mountains, and Charleston Peak. From Tingley et al. (2001), Geologic Tours in the Las Vegas Area.

Some debates consider whether sediments are tills or debrisﬂow sediments containing reworked till; a still-unresolved example is the southern Alberta “Bow Valley till” of Rutter (1972; see also Mandryk and Rutter, 1990), which was reinterpreted as debris ﬂow sediment by Eyles et al. (1988, 1990). Owen (1991, 1993) attempted to distinguish properties of Quaternary tills and debris ﬂows in the Karakorum, but found distinction difﬁcult because of reworking of tills in the ﬂows. In particular, there was little to distinguish between debris ﬂows and supraglacial tills. Subglacial tills were more likely to be overconsolidated and compact, to exhibit some edge-crushing of grains, and to contain coherent systems of microshears than debris ﬂow sediments, but liquid limits, clay mineralogy, and grain shapes were similar between the two sediment types (Owen, 1991). Given the potential uncertainties in interpretation of glacial environments, various authors have described criteria, which tend

to focus on sedimentary records, presumably because erosional records are more easily lost with time. Benn and Evans (1998), for example, describe various kinds of till and associated sediments, and also provide a useful summary of glacial landforms. Boulton (1978) and Eyles (1993) believe bullet-shaped clasts are one of the most useful indicators of glacial transport. Hambrey and Harland (1981) compiled a list of evidence for terrestrial glaciation (shown here in Table 1), while Goldthwait (1971) listed characteristics of glacial till (Table 2). Complicating the situation are (1) the fact that any one of these criteria in isolation can be found in non-glacial environments, and (2) disputes over the value of some of these criteria. For example, Spenceley (2001, 2003) and Gore and Taylor (2003) argue over glacial versus non-glacial origin of bedrock grooves and striations; Bennett et al. (1999) suggest that clast fabric offers little quantitative support in the interpretation of

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Osborn et al. TABLE 1. INDICATORS OF TERRESTRIAL GLACIATION ACCORDING TO HAMBREY AND HARLAND (1981) Abraded bedrock surfaces Stone-rich beds Depositional fossil landforms of distinct form (e.g., drumlins, moraines) Variable lithologies of stones Stones with a wide range of shapes (but especially subrounded and subangular) Striated and faceted stones Clay-sized particles (rock flour) Clay or calcareous “eggshells” (thin coatings on larger stones) Fragile stones (e.g., shale fragments) mixed with quartzose stones Quartz grain textures Chattermark trails on garnets

TABLE 2. CHARACTERISTICS OF GLACIAL TILL ACCORDING TO GOLDTHWAIT (1971) Poorly sorted, often with clasts of many sizes in a variable finer matrix Tends to be massive in structure, without smooth lamination or graded bedding Composed of mixed minerals and rock types, some of which are far-traveled Generally contains some striated stones and micro-striated grains May have a common orientation of elongated particles (fabric) May be quite compact May rest upon striated rock or sediment basement Component clasts are predominantly sub-angular

glacigenic sediments; Whalley and Krinsley (1974) suggest that grain micro-textures may not be able to distinguish glacial from some non-glacial environments; Eyles (1993) claims that supposedly unique “glacial textures” are found on grains subjected to repeated breakage, regardless of sedimentary environment. THE SPRING MOUNTAINS AND VICINITY Modern Regional Climate The Mojave Desert and southern Great Basin are located in the U.S. desert southwest, and contain a diverse and unique geologic and geographic landscape. North-south–trending mountain ranges may reach elevations above 12,000 ft (3660 m) and are separated by basins with elevations as low as 282 ft (86 m) below sea level at Death Valley. The region today comprises the driest area in the continental United States, but it hosted large pluvial lakes in the recent past that document signiﬁcant variations in temperature and effective moisture. The climate of the Mojave Desert and southern Great Basin varies by elevation and distance from the coasts. Mean annual temperatures reﬂect the continentality of the region. Rainfall and snowfall are more abundant at high elevations, which, coupled with decreased temperatures and evaporation with elevation, results in an altitudinal gradient in effective moisture. The

mean annual temperature at McCarran International Airport in Las Vegas (659 m [2160 ft] a.s.l.), is 19.5 °C (67 °F), and mean annual precipitation is 107.4 mm (4.2 in) (data available at www. worldclimate.com). In contrast, rainfall increases to 307.5 mm (12.1 in) at the town of Mt. Charleston (2288 m, 7506 ft), reﬂecting the orographic effect over the high Spring Mountains. Temperature at Kyle Canyon (2195 m, 7200 ft) averages 11.0 °C (52 °F) (data available at http://pnwpest.org/SW/NV/index.html). Much groundwater recharge to basins and stream runoff is the result of winter snow melt in the spring season, but summer monsoon rainfall may also recharge groundwater (Smith et al., 1992), particularly in higher forested areas. The region is underlain by the Paleozoic carbonate aquifer, a large-scale groundwater ﬂow system encompassing much of Nevada, including Yucca Mountain and the Nevada Test Site, and terminating at Death Valley, California. Rainfall in the southern Great Basin is produced under two distinct dominant climatic regimes: (1) midlatitude frontal Paciﬁc cyclones, and (2) the North American Monsoon (NAM). The NAM results in convection-dominated summer rainfall, whereas the Paciﬁc cyclones produce frontal rainfall during winter months. The moisture paths differ between these two systems (Friedman et al., 2002). Paciﬁc cyclonic systems are borne by the westerlies and deliver winter moisture to the southern Great Basin. The efﬁciency of the orographic rainout results in a distinct rain shadow in the lee of the north-south mountain ranges, including the Sierra Nevada and White Mountains, which results in the arid to semiarid climate of the Mojave Desert and southern Great Basin. Summer precipitation is more subject to evaporation than winter precipitation (Smith et al., 1992), and is thus less likely to contribute to groundwater recharge than winter precipitation. The effect of the El Niño–Southern Oscillation on rainfall in the southern Great Basin is strong. During El Niño events, a southward displacement of the jet stream results in the delivery of anomalously high precipitation to southern Nevada (Sharp, 2003). In contrast, La Niña phases are associated with anomalously low precipitation. Southward shifts in storm tracks have also been suggested for the wetter pluvial conditions of the late Quaternary, although the role of El Niño in such a shift is poorly known. Late Quaternary Paleoclimate in the Mojave Desert Region The paleoclimatic history of the Mojave Desert and surrounding regions is known in broad detail for the past 0.5 million years (Sharp, 2003), based on numerous studies that utilized lacustrine (Enzel et al., 2003), marsh (Quade et al., 2003), and ﬂuvial sediments (Cox et al., 2003), saturated zone calcite deposits (Coplen et al., 1994; Winograd et al., 1988), eolian deposits, packrat middens (Jennings and Elliott-Fisk, 1993), tree rings (LaMarche, 1974, 1978), and glacial records from the Sierra Nevada (Clark et al., 2003). The geologic evidence available suggests that now-dry lake basins have ﬁlled with water on multi-millennial glacial to interglacial time scales (Enzel et al., 2003), in concert with

Pleistocene glaciation in the Spring Mountains of Nevada pronounced vegetation changes (Coplen et al., 1994). In contrast to the Northern Great Basin, pluvial lakes in southern Nevada were smaller, or better characterized as paludal marsh and spring deposits, and they typically lack former shoreline geomorphology. The three longest paleoclimate records for the area are the Devils Hole calcite, sediments recovered from Owens Lake, east of the Sierra Nevada near Lone Pine, California, and a drill core from the Badwater playa, Death Valley, California. The Devils Hole record (Fig. 3) is from a uranium-series dated calcite deposit, where δ18O variations are most strongly linked to temperature variations in the Great Basin. The original record spans the interval from ca. 500 ka to 60 ka, and the record has been recently extended to 4.5 ka by analysis of additional material (Winograd et al., 1992, 2006). The record shows cyclical variations in δ18O that yield low values during cold glacial times and high values during interglacial times, which suggests a dominant temperature effect in the precipitation arriving at Devils Hole. The recharge area for Devils Hole includes the Spring Mountains, so the climate signal from there should also reﬂect climate conditions in the Spring Mountains. For the late Quaternary, coldest conditions are recorded at ca. 65 ka and 30 ka. Warming to the Holocene commenced at ca. 25 ka, and was complete ca. 18 ka. At Owens Lake, sediments record cyclical wet periods during glacial stadials, alternating with dry periods during interglacials (Benson et al., 1996). The lake basin constantly overﬂowed between 52.5 and 23.5 ka, partly as a result of meltwater delivery from glaciers in the high Sierra Nevada, and partly due to a wetter climate. The timing of dry periods around Owens Lake, and apparently contemporaneous glacial advances (dated with 36Cl cosmogenic nuclides) suggest a link to Heinrich ice-rafting events in the North Atlantic Ocean (Benson et al., 1996; Phillips et al., 1996). The revised ages of four separate Tioga advances in the eastern Sierra Nevada (Bloody Canyon and others) range in age from 25 to 14 ka (Kaufmann et al., 2003). The Mono Basin and Tahoe moraines are considered to date to Marine Isotope stages 3–6, although controversy remains on their true ages (Kaufmann et al., 2003). The Tioga moraines represent the last glacial moraines and are well-expressed throughout the Sierra Nevada. The Tenaya moraines may represent an early Tioga advance, or may be a discrete advance; further chronology is required to evaluate these possibilities (Kaufmann et al., 2003). Following the last glacial period, a late-glacial Recess Peak advance occurred ca. 14.2–13.1 ka and was followed by small glaciers that formed during the Neoglacial period in and near cirques. The Death Valley sediment core reveals alternating salts and muds, representing dry and less-dry conditions respectively (Lowenstein et al., 1999) occurring on an ~100 ka timescale. Wet periods sufﬁcient to sustain perennial lakes were present at ca. 120–186 ka, and from 10 to 35 ka. Of these two wet periods, the penultimate glacial period was characterized by wetter conditions. In the Las Vegas Valley area, spring deposits at Corn Creek Flat, just east of the Spring Mountains, were assigned to the Las Vegas Formation by Longwell et al. (1965). They date to the full

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Figure 3. Paleoclimatic summary for southern Nevada, showing the Devils Hole calcite δ18O record (continuous line and points; higher values indicate warmer temperatures); periods of perennial lakes in Death Valley (dashed horizontal lines; from Lowenstein et al., 1999); the approximate timing of deposition of the Las Vegas Formation (dashed and solid line indicating intermittent and continuous wet periods; from Quade et al., 2003), and the glacial stillstands in the Sierra Nevada (from Kaufmann et al., 2003, based partly on 36Cl boulder dating). Figure modiﬁed from Winograd et al., 2006.

glacial (25–14 ka), and less abundant late glacial (14–8 ka) periods (Quade et al., 1995). These deposits are visible from U.S. Highway 95 as white-colored silts, and represent a former phreatophyte ﬂat in a wet meadow environment. From these data, the water table level appears to have been ~25–40 m higher than today. Black mat deposits associated with fossil springs are present in southern Nevada, and were primarily formed during periods of enhanced effective moisture between ca. 11.8 and 6.3 ka (Quade et al., 1998), and may be related to wetter conditions during the Younger Dryas event. Geology and Geography of the Spring Mountains The Spring Mountains are a NW-SE–trending fault block mountain in the Basin and Range Province of southern Nevada (Fig. 1). The mountains are delimited to the northeast by the Las Vegas Valley Shear Zone, and to the southwest by the Pahrump Valley. The geology of the Spring Mountains is dominated by thrust-faulted Cambrian through Pennsylvanian sedimentary rocks overlying the Jurassic Aztec Sandstone, and some Permian and Triassic rocks (Fig. 4). The Spring Mountains are well known for the beautiful exposures of the Jurassic Aztec Sandstone (known as the Navajo Sandstone in other regions of the

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Figure 4. Generalized stratigraphy for the Spring Mountains of Southern Nevada. From Tingley et al. (2001), Geologic Tours in the Las Vegas Area.

Figure 5. Altitudinal variations in precipitation and temperature inﬂuence the life zones in the Spring Mountains. From Tingley et al. (2001), Geologic Tours in the Las Vegas Area.

Pleistocene glaciation in the Spring Mountains of Nevada American southwest) in the southern portion at the Red Rock Canyon National Conservation Area. The eolian Aztec Sandstone is a world-famous rock climbing locality because of the numerous hand holds and competence of the rock. The Aztec sandstone is overlain by Cambrian Bonanza King limestone and dolomite in the compound Keystone thrust system. The eastern foot of the range lies at an elevation of ~4000 ft (1220 m) a.s.l. while the high point, Charleston Peak, rises to11,918 ft a.s.l. (3634 m). Much of the range drains to the northeast, into Las Vegas Valley, via several canyons, of which the two largest, Kyle and Lee, are accessible by paved road. The canyons contain terraced and carbonate-cemented alluvial ﬁlls which reach elevations of ~8000 ft (2440 m) near the canyon heads (Dolliver, 1968) and extend beyond the canyon mouths as coalescing alluvial fans. The coalesced fans form an alluvial apron that descends to an elevation of ~2500 ft (760 m) in the center of Las Vegas Valley (Lattman and Simonberg, 1971). Bedrock slopes in the higher elevations of the range are developed in resistant Paleozoic limestones and are generally steep. Bases of slopes are veneered with extensive colluvial deposits which locally head close to the crest of the range and extend downward to grade into or interﬁnger with alluvial fan sediments. Rainfall in the area ranges from ~20 in (50 cm) per year near the crest of the range to ~4 in (10 cm) per year in the Las Vegas Valley (Longwell et al., 1965). Considerable snow accumulates in the higher elevations of the range, allowing a ski resort in Lee Canyon to serve the local residents in winter. Severe spring ﬂoods in the canyons, which may spread out across alluvial aprons distally, were common (Lattman and Simonberg, 1971), but are increasingly intercepted by engineered detention basins. The high vertical relief in the Spring Mountains results in numerous life zones that are related to the decreasing temperature and increasing precipitation with altitude (Fig. 5). INTERPRETATION OF GLACIATION IN THE SPRING MOUNTAINS Glaciation in the Basin and Range Province Many of the fault-block ranges in the Basin and Range Province have been glaciated. The southernmost glaciated range on the border of the Basin and Range Province is the San Gorgonio Mountains in the Transverse Ranges of southern California (Sharp et al., 1959; Owen et al., 2003), which lie ~250 km of latitude south of the Spring Mountains. The high point of the range is 3506 m (11,499 ft); the longest Pleistocene glaciers were 2.5 km long. Owen et al. (2003) attribute the unusually low paleo-ELA for that latitude to substantially increased winter precipitation under a maritime inﬂuence, coupled with increased presence of the jet stream over southern California in the Pleistocene. Osborn and Bevis (2001) listed 40 glaciated ranges other than the Spring Mountains either entirely within or lying on the boundary of the Great Basin, the large region of internal drainage within the Basin and Range Province. This list includes the

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Sierra Nevada and Wasatch Ranges, generally regarded as the western and eastern bulwarks of the Basin and Range Province. The Spring Mountains are on the border of the Great Basin, with the east ﬂank draining into the Colorado River system. The closest glaciated ranges to the Spring Mountains listed by Osborn and Bevis (2001) are the southern Sierra Nevada to the west, the White Mountains 250 km to the northwest (high point 14,245 ft or 4343 m; glaciers to 6 km long); the Monitor Range 250 km to the north-northwest (high point 10,888 ft or 3320 m; glaciers to one km long); and the Grant Range 170 km to the north (high point 11,300 ft or 3445 m; glacier 0.5 km long). Glaciation in the Spring Mountains: History of the Concept The high peaks of the Spring Mountains have long been considered as possible candidates for Quaternary glaciation, because of their high altitude and suggestive geomorphology. Eliot Blackwelder, the ﬁrst to review glaciation in the Basin and Range Province, made no mention of the Spring Mountains in his original paper (1931); in his supplementary paper (1934) he ﬁgured the “Spring Mountain Range” was too far south to be glaciated during the Tioga glaciation, but that “The somewhat excavated heads of the canyons surrounding the highest peak afford a suggestion of glaciation during the Tahoe stage” (p. 221). In his table of “Preliminary Data on Glaciated Areas in Western United States” Flint (1947, 1957) included Spring Mountain [sic] as the site of a cirque glacier, on the basis of an unpublished manuscript or communication with C.R. Longwell. But in Flint’s 1971 edition, there is no mention of the Spring Mountains. No one seems to know the identity or whereabouts of Longwell’s manuscript, if indeed that’s what it was, and Longwell et al. made no mention of glacial features in their geologic study of Clark County (1965). Other Spring Mountains overviews besides Longwell et al. (1965) are Burchﬁel (1974) and Page et al. (2005); no glacial features are interpreted in these works. Of geomorphic studies, Dolliver (1968) maps no glacial deposits or features, while Lattman and Simonberg (1971) state “There is no evidence of glacial deposition, and the only glacial effects seen are possible small cirques in some of the highest parts of the range.” Piegat’s (1980) thesis on glacial geology in Nevada notes “On Charleston Peak there are no identiﬁable landforms resulting from glacial erosion or deposition. However, at the head of Kyle Canyon on the east ﬂank are three or four features which may be badly degraded cirques” (p. 74). But Piegat ends up suggesting that even the highest point of the Spring Mountains was too low to intersect the Pleistocene snowline (p. 75). Nick Saines found an interesting diamict below Big Falls in the upper Kyle Canyon drainage in the early 1990s and showed it to Rick Orndorff and John van Hoesen, then at the University of Nevada, Las Vegas, in the late 1990s. This group regarded the diamict as a till, mainly on the basis of striated clasts and the interpretation of the landform as a moraine, and liked the preexisting idea of degraded cirques. They published abstracts (van Hoesen

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Figure 6. Oblique Google Earth image looking to the west-southwest of the highest portion of the Spring Mountains, including Charleston Peak and upper Kyle Canyon, with sites mentioned in the ﬁeld trip guide indicated. Vertical elevation difference between Charleston Peak and the Kyle Canyon diamict is 1100 m.

et al., 2000; van Hoesen and Orndorff, 2000) and two papers (van Hoesen and Orndorff, 2003; Orndorff et al., 2003) promoting the concept of glaciation in the Spring Mountains, prior to late Wisconsinan time, at one point claiming to have ended the longstanding debate (Orndorff et al., 2003, p. 1) The idea of glaciation in the desert attracted some attention in the Las Vegas media. Orndorff et al. (2003) is available in the GSA Data Repository1. Osborn and Bevis (2001) published a review of glaciation in the Great Basin and listed the Spring Mountains with a question mark, questioned the combination of presence of old glacial deposits and absence of late Wisconsinan deposits, and noted that “Further work will be required to establish the glacial (or nonglacial) origin of the deposit.” Most recently, Osborn et al. (2008; see Appendix 1) point out problems with the interpretations of Orndorff et al. (2003). The Debate In this discussion, OVS will refer to Orndorff, van Hoesen, and Saines, who published the glacial interpretation (Orndorff et al., 2003) and OCL will refer to Osborn, Clark, and Lachniet, whose opinions are brieﬂy summarized in Osborn et al. (2008). 1 GSA Data Repository item 2008082, Orndorff et al., 2003, Implication of new evidence for Late Quaternary Glaciation in the Spring Mountains, Southern Nevada (Journal of the Arizona-Nevada Academy of Science, v. 36, p. 37–45), is available at www.geosociety.org/pubs/ft2008.htm, or on request from editing@geosociety. org, Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.

OVS base their glaciation scenario on interpretation of (a) valley head geomorphology of upper Kyle Canyon as glacial in origin, and of (b) two particular exposures of diamict as till. In their view, a cirque and a short length of U-shaped valley lie in a tributary above a waterfall known as Big Falls, and other “cirquelike features” lie at the head of Kyle Canyon, at a lower elevation than that of Big Falls (Fig. 6). Because there are no smallscale glacial-erosional features on the headwalls, they regard the cirques to be Tahoe or Mono Basin (of Sierran terminology) in age, and conclude that there was no late Wisconsinan glaciation in the range. The knickpoint at Big Falls leads OVS to refer to the valley head above as a “hanging valley,” presumably to further establish the glacial origin of the valley. OCL regard the “cirque” above Big Falls (Fig. 6) as an ordinary valley head developed on cliff-and-bench topography in horizontal, stratiﬁed rocks. In their view, (a) the upper “U-shaped valley” of OVS is a consequence of the erosional resistance of the upper unit of the Mississippian Monte Cristo Formation: that unit creates the bench on which lies the upper valley bottom; (b) the upper valley “hangs” because the drainage of the upper valley eventually cascades down the Monte Cristo cliff to the less resistant rocks below; and (c) the other, lower-elevation “cirque-like features” at the head of Kyle Canyon (e.g., “putative cirque” on right side of Figure 6) are merely embayments, along drainage lines, in the resistant Monte Cristo cliff. OCL regard the presence of lower-elevation cirques at the head of Kyle Canyon, without cirques above at the actual valley

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Figure 7. Big Falls wash: Exposures marked here as “Kyle Canyon diamict” and “Hanging diamict” are interpreted as till by Orndorff et al. (2003). Vertical elevation difference between top of Big Falls and Kyle Canyon diamict is ~200 m.

Figure 8. Part of the Kyle Canyon diamict. Middle part dips down to the left; diamict in lower right is interpreted as till by Orndorff et al. (2003). Arrow points to person at lower left for scale.

heads, as very unlikely, and perhaps even inconsistent with the deﬁnition of “cirque.” They see no evidence of glacial erosion, at any scale, in upper Kyle Canyon. OVS interpret a small ridge at the side of a wash, at the mouth of the Big Falls tributary, as an eroded remnant of a terminal moraine constructed by a glacier that ﬂowed down the Big Falls tributary (“Kyle Canyon diamict”; Figs. 6 and 7). OCL regard the moraine interpretation as untenable: the ridge runs parallel to the valley so it has the wrong orientation to be an end moraine; meanwhile it lies at the very bottom of the valley and so cannot be a lateral moraine.

OVS interpret the diamict exposed in the ridge (Fig. 8) to be till capped by 2.2 m of well-sorted and stratiﬁed glacioﬂuvial sediment; the central portion of the deposit shows “upward trend toward vague stratiﬁcation,” perhaps due to reworking of the till by mass wasting or erosion. The till interpretation is based on striated clasts that are additionally described as polished and faceted. Orientation analysis was done on 15 striated clasts and a preferred orientation along long axes of clasts was noted. OVS also describe micro-textures on limestone clasts that they imply to be evidence of glacial abrasion, although they cite no literature in support of the implication.

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OCL do not agree with the basic description of sediments within the exposure, but instead see stratiﬁed diamict dipping at ~30° in the downstream part of the exposure, capped by a meter or two of local slope colluvium; going upstream (which is going down section because of the dips of the layers) the stratiﬁed diamict gradually changes into very-vaguely to unstratiﬁed diamict. OCL see no demarcation or contact between the stratiﬁed and unstratiﬁed parts of the sediment and believe it to be basically the same material. They regard lack of bimodality of grain sizes, angularity of most of the clasts, stratiﬁcation of much of the sediment, dip of the stratiﬁcation away from bedrock walls and toward the valley axis, and location of the sediment in a valley bottom where ongoing mass movement and ﬂashy alluviation are evident, to indicate that the sediment is probably not till, but rather colluvium possibly mixed with some alluvium. Furthermore, they believe no ridge of till from a pre-Wisconsinan glaciation would survive a location in a valley bottom, along a wash, where ﬂash ﬂoods and debris ﬂows are not uncommon; a relatively great age of the deposit is also contradicted by only minor soil formation and caliche development, shown by stage II calcite pendants on limestone clasts near the surface of the downstream portion of the ridge. OCL think the striation-orientation analysis of OVS was on too small a population to necessarily be valid, and are not aware that anyone has shown that the micro-textures described by OVS cannot occur in non-glacial environments. OVS describe a second exposure of till along the east-facing slope above Big Falls wash, a few hundred meters upstream from the ﬁrst exposure (“Hanging diamict” in Fig. 7; Fig. 9). OCL regard that material as slope colluvium.

OCL agree with OVS that many clasts in the diamicts at both exposures are striated, and that some of the striations occur on fairly smooth faces of the clasts (Fig. 10). To proponents of glaciation here, the striations are probably the strongest evidence of glacial inﬂuence. Hence they are of key importance in the debate over origin of the deposits. The Meaning of Striations Striations (also known as “striae”) are “among the most common features of glacial erosion” (Hambrey, 1994) and one of the most widely recognized means of establishing a glacial environment (e.g., Hambrey and Harland, 1981; Aitken, 1991; see also Atkins, 2003). But there is inherent ambiguity in the use of striations. Hambrey (1994) notes that striations are insufﬁcient in themselves as glacial indicators since other, non-glacial agencies can give rise to them. Several authors in the twentieth century reported non-glacial striations and some cautioned against an automatic glacial interpretation (e.g., Hovey, 1909; Wentworth, 1928; Dyson, 1937; McLennan, 1971; Zamoruev, 1974). Schermerhorn (1974) observed that cautions by such authors did not have much inﬂuence on the wider community. Eyles (1993) referred to means of creating striated clasts and listed examples where deposits originally identiﬁed as “glacial” were later shown to be formed in non-glacial environments. Atkins (2003; p. i) stated “…there are many non-glacial processes that can produce striae. They have been sporadically documented in the geological literature but have failed to make a lasting impression on the wider Earth Sciences community.”

Figure 9. Part of the Hanging diamict; interpreted as till by Orndorff et al. (2003). Lower edge of frame is ~4 m wide.

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Figure 10. (A, B, and C) Striated clasts from the Kyle Canyon diamict. (D, E., and F) Striated clasts from the Hanging diamict. Coin and pen for scale.

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Figure 11. Striated clasts from non-glacial environments from (A) debris ﬂow on Mount Adams, Washington, (B and C) channel of Cottonwood Wash, Death Valley, (D) debris ﬂow on Hudson Bay Mountain, British Columbia, (E) rockfall talus on Hudson Bay Mountain, British Columbia, and (F) bedload of the Owyhee River, Oregon. Coins and carabiner for scale.

Pleistocene glaciation in the Spring Mountains of Nevada Striated clasts in mass moved sediments are reported by, among others: Palmer and Neall (1989; debris avalanche); Sundell (1985; debris ﬂow); Jensen and Wulff-Pedersen (1996; debris ﬂow); Harrington (1971; mudﬂow); Blackwelder (1930; mudﬂows); Van Houten (1957; mudﬂow); Winterer and von der Borch (1968; mudﬂow); Scott (1988; lahar); and Sparks et al. (1997; lahar). Many of the deposits described in these studies were originally interpreted as tills. Non-glacial striated clasts observed by Osborn are shown in Figure 11. Atkins (2003) attempted to ﬁnd characteristics of striations that could serve as discriminators of different environments. He compared striations from temperate, polythermal, and cold glacial environments with those from mass movement and tectonic environments. He concluded that glacial striations tend to be subparallel to clast long axes and show a high density on individual surfaces, whereas those of non-glacial origin tend to show a lower density of slightly shorter, wider striations and show either weak orientation or none at all. These are tendencies or means; his data show overlap between examples from different environments for any particular attribute. Furthermore, he examined clasts from only two cases of mass movement, one rockfall deposit and one debris avalanche. Nevertheless, if large sample populations are

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considered, his conclusions offer some help in distinguishing glacial from mass movement striations, the two most likely candidates in the Spring Mountains. ROAD LOG This ﬁeld trip will proceed up U.S. Highway 95 to Kyle Canyon Road and then up that road, to focus on landforms and sediments at the head of the canyon. The road log also includes the road between Kyle and Lee canyons and the Lee Canyon Road, for independent interested parties that may wish to do a loop trip (Fig. 12). The road log begins at the parking lot on the east side of the Student Union building at the University of Nevada, Las Vegas. Mileage 0.0 mi 0.5 2.9 7.9

Description Leave Student Union, drive south on Maryland Parkway. Turn west onto Tropicana Ave. Take on-ramp to Interstate 15 North. Take on-ramp to U.S. 95 North (actually west at this point).

Figure 12. Shaded relief map of Kyle and Lee Canyons east of Charleston Peak. From Tingley et al. (2001), Geologic Tours in the Las Vegas Area.

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Osborn et al. From the overpass connecting I-15 to U.S. 95, the mountains surrounding Las Vegas can be seen. Frenchman and Sunrise Mountains are to the east, the Las Vegas and Sheep ranges are to the north, and the Spring Mountains are to the west and northwest. To left (south) is Las Vegas Springs Preserve. The 180-acre Las Vegas Springs Preserve opened in 2007. The Las Vegas Springs used to ﬂow at surface, the groundwater coming up along faults, with the recharge area in the Spring Mountains to the west. The springs created a natural oasis in the desert, with evidence of Native-American habitation going back thousands of years. The springs were an important watering hole on the Old Spanish Trail, providing grassy meadows or las vegas for the animals driven along the trail. The springs were visited by explorer Captain John C. Fremont and Kit Carson in 1844, and provided the water for the Mormon settlement of Las Vegas in 1855–1857. The springs stopped ﬂowing in 1962 as the potentiometric head dropped due to groundwater pumpage. The Preserve includes walking trails, museums, a theater, desert demonstration gardens, and a conference center. The new Nevada State Museum is under construction here. To right (east), in the lowest part of valley, is the Las Vegas Formation. According to Longwell et al. (1965) the Las Vegas Formation is made up of light-colored thinly layered deposits of clay and silt, containing abundant shells of freshwater snails and other mollusks, suggesting a spring-fed paludal and/or lacustrine environment. Fossils of extinct mammoths, ground sloths, American Lion, camels, llamas, bison, and horses, ranging in age from 40,000 to ca. 11,000 ka, have been found in these deposits. Efforts are under way to preserve a large portion of Upper Las Vegas Wash as a paleontological preserve. Junction with State Route 157 (Kyle Canyon Road). Turn left. Highway 157 climbs the Kyle Canyon alluvial fan, one of the large fans ﬂanking the east side of the Spring Mountains. It was built by ephemeral streams, most notably Kyle Canyon Wash and Harris Springs Wash (Sowers, 1988). There is no evidence that external drainage ever ﬂowed through Kyle Canyon; all material in the fan is indigenous to the present drainage basin (Dolliver, 1968). The fan is the surface expression of a variably thick sedimentary ﬁll derived from weathering and mass wasting of the mountain sides; in the case of Kyle Canyon the sediment buries a steep-sided V-shaped bedrock valley, the bottom of which at one point was at least 740 ft (225 m) deep, prior to dissection

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of the fan (Dolliver, 1968). Dolliver suggests that the sedimentary ﬁlling was initiated in late Tertiary or early Pleistocene time, during a period of increased runoff and cooler climate. The total area of the drainage basin is 450 km2, of which 220 km2 consists of alluvial deposits (Sowers, 1988). According to Dolliver (1968), the material in the Kyle Canyon fan consists of colluvium, waterworked colluvium, and alluvium, characterized by increased stratiﬁcation, sorting, and clast roundness toward the alluvial end of the spectrum. Colluvium close to the source is described as unstratiﬁed and unsorted with great variation in grain size. Dolliver notes great lateral and vertical variation in the sediment. Sowers (1988) deﬁne six “alluvial units”; they all consist mainly of coarse carbonate detritus but have varying amounts of siltstone and sandstone and one has beds that are tilted slightly relative to the others. Tingley et al. (2001) note that “the crudely deﬁned and irregular layering of the boulders, cobbles, pebbles, and sand displayed by this fan deposit is typical of material carried by debris ﬂows or mudﬂows” (p. 39). Sometime following canyon ﬁlling, excavation of the ﬁll began. Dolliver (1968) interprets three periods of incision and two major periods of valley widening, which resulted in three major levels of modern fan surface. Sowers (1988) identiﬁes four geomorphic surfaces on the fan. According to the latter author the evolution of the fan surface consisted generally of erosion (i.e., trenching) of older surfaces in the upper elevations, and simultaneous burial of older surfaces in the lower elevations (i.e., progradation of the fan). The fan surfaces all are characterized by accumulations of pedogenic carbonate, ranging from clast coatings and partial void ﬁllings in the younger surfaces to calcretes in the older surfaces. Calcrete paleomagnetism indicates that the oldest surface is >730,000 years old (Sowers, 1985). The white scar of a bedding-controlled rock slide can be seen high on the ridge south of Kyle Canyon (Fig. 13). STOP 1: Diamict in Kyle Canyon fan (200 m downstream of The Castle) At this stop we will examine some of the variety of sediment types in the fan, some of which in terms of process could be distantly related to diamicts we will see at the head of the canyon. The exposures here are in the north wall of the active (although almost always dry) channel of Kyle Canyon Wash. The sediments are stratiﬁed overall but include 0.5–1 m intervals of non-stratiﬁed or very faintly stratiﬁed gravel, intervals of moderately wellstratiﬁed gravel, and irregular lenses of sand. There

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Figure 13. Block diagram illustrating rocksliding along bedding planes in the tilted carbonate strata of the Spring Mountains. From Tingley et al. (2001), Geologic Tours in the Las Vegas Area.

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is a spectrum of openwork gravel to sandy-matrix gravel. The sediments are cemented to varying degrees by calcium carbonate. Worthy of discussion here are the meaning of “diamict,” the distinction between and interpretation of debris-ﬂow fanglomerates and alluvial fanglomerates, and meaning of angularity versus rounding of gravel clasts. Junction with Highway 158. Continue straight ahead on Highway 157. Upstream of here, Kyle Canyon is bounded by mountain ridges. Parts of the canyon have a U-shaped transverse cross section similar to that of a glaciated valley (Fig. 14), but the “U” is developed on valley ﬁll and colluvial aprons described by Dolliver (1968), and it has no connection to glacial erosion. Leave the highway at the start of the hairpin turn, onto a secondary paved road that leads through a collection of cabins and summer homes, part of the community of Mt. Charleston (a misnomer since the summit of the range is actually called Charleston Peak). The upper levels of the community are built on terraces of the Kyle Canyon fanglomerates, which before dissection rose to a level at least 40 m above the road. STOP 2: Recent avalanche, illustrating one type of mass movement process in the range. This snow avalanche occurred in January of 2005. The mass of snow, rock, and broken trees blocked the road. After the road was cleared the snow in the avalanche deposit still had not melted by the summer. Where the paved road bends to the right (toward Trail Canyon) keep straight ahead on a gravel road.

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Figure 14. Looking downstream in Kyle Canyon from Cathedral Rock. Parts of the canyon have a cross-section similar to that of a glaciated valley.

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Go to the end of the road into the Mary Jane Falls trail parking lot and park. The parking lot and the trail to Mary Jane Falls are in the bottom of upper Kyle Canyon Wash, which is incised into the Kyle Canyon fanglomerates. Projection of the top of the highest remaining fan surface, just upstream of the gravel-road turnoff from the paved road, indicates that the original fan surface extended high above the parking lot. Bedrock strata here are approximately horizontal. To the north and south of the parking lot, and wrapping around the very head of Kyle Canyon 2 km to the west, is an obvious, very steep cliff developed in the upper Monte Cristo Formation of Mississippian age. Big Falls, on the south side of the canyon, and Mary Jane Falls, on the north side, both are consequences of that resistant cliff-former. The lower-gradient slopes extending from the cliff base down to the wash are a lower unit of Monte Cristo. The lower-gradient slopes above the top of the cliff, which gradually steepen upward, are in the Pennsylvanian-Permian Bird Spring Formation. STOP 3: Hike (about one mile) to Big Falls wash, site of “Kyle Canyon till” and “Hanging till” of Orndorff et al. (2003). Proceed up the Mary Jane Falls trail to where it leaves the wash bottom and starts to switchback up the north canyon slope (~20 min). Leave the trail and proceed southwestward across the gravel ﬂats on the remains of an old gravel road. Note the low-relief hummocks on the ﬂats, which are most likely degraded remnants of old debris-ﬂow lobes and levees, and/or avalanche deposits. After a short distance, the trees begin to thin

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Osborn et al. out and a great swath of avalanche-ﬂattened forest can be seen to the northwest, in uppermost Kyle Canyon. The hummocks and downed trees are indicative of an active mass movement environment in the area. The landform containing the putative Kyle Canyon till is located at the mouth of Big Falls wash, where it joins the main wash from uppermost Kyle Canyon, a 15 min walk from the Mary Jane Falls trail (Figs. 6 and 7). We will ﬁrst consider the landform from a viewpoint across the main wash and then from across Big Falls wash (i.e., northwest side of wash and slightly upstream). From this latter point, one gets a good view of the Monte Cristo cliff, Big Falls wash and the general site of Big Falls itself, uppermost Kyle Canyon and its ﬂattened forest, and Mary Jane Falls. The landform in question is a ridge that follows, in fact forms in part, the south wall of Big Falls wash. A dip in the ridge crest separates an upstream part of the ridge with poorly exposed sediments from a downstream part with well exposed sediments. The ridge is considered by Orndorff et al. (2003) to be a lateral moraine of a pre-Wisconsinan glacier that ﬂowed down Big Falls wash. Other possibilities suggested by various observers are that the ridge (a) is a lateral or terminal moraine of a glacier that ﬂowed from the very head of Kyle Canyon, (b) is a large debris ﬂow levee, (c) is a protalus rampart, or (d) was not constructed as a ridge and only looks like one now because incision of the wash removed part of the

original landform. These hypotheses will be considered during a circuit of the landform. Cross Big Falls wash and climb up the southeast bank, through the dip in the ridge crest, to the red snow marker. Here can be seen the linear, gullylike depression behind the ridge crest, which parallels the crest and the wash. A short walk to the rock outcrop at the head of this depression shows that the depression has no drainage source, and on that basis does not appear to be erosional. If the ridge is a pre-Wisconsinan lateral moraine along Big Falls wash, the glacier at this point must have been very narrow and very thin. Apart from any questions of sedimentology, it is problematic that a moraine that old would survive the dynamic and ﬂashy alluvial and mass movement environment that characterizes this locality. Furthermore, the inceptisol on the crest of the ridge may not be consistent with a pre-Wisconsinan age of the sediment. A levee or protalus origin of the ridge would allow for a young age and still explain the depression behind the ridge crest. It is possible that the parts of the ridge up- and downstream of the dip in the crest may be of different origin. Walk down the depression to the downstream end of the ridge and circle around to the good exposure visible from the wash (Fig. 8). The downstream part of the exposure consists of poorly stratiﬁed diamict dipping ~30° NW, with some lenses of openwork gravel (Fig. 15), and overlain by a meter of two of colluvium that parallels the local

Figure 15. Dipping openwork gravels in the Kyle Canyon diamict.

Pleistocene glaciation in the Spring Mountains of Nevada slope. Upstream, the vague stratiﬁcation is mostly or entirely lost, depending on the observer. The sediment is unsorted except for a few pockets of vaguely sorted sand. There is no visible demarcation or contact between the stratiﬁed and non-stratiﬁed parts of the diamict; Occam’s Razor suggests that both parts were deposited in the same environment. Pebbles and cobbles are angular to subrounded, and many are striated in both the stratiﬁed and non-stratiﬁed parts of the sediment. The dipping deposits on the downstream end are interpreted as glacioﬂuvial sediment by OVS. But most of the deposit appears to be too unsorted to be ﬂuvial in origin; furthermore, the up-dip projection of the beds does not follow Big Falls wash but rather leads up toward the bedrock cliff, so it is not clear where the glacial ice would have been that provided those sediments. The alternative is that the dipping beds are some form of colluvium, possibly debris ﬂow sediments. This possibility was independently suggested in a Web site by Abe van Luik (http://www.thoughtsandplaces.org). In the protalus hypothesis, the colluvium would have collected at the bottom of a snowbank that lined the base of the bedrock tower rising above the ridge. However, the beds seem to project up dip to slightly south of the tower, and it is not clear exactly where the colluvium would have been coming from. The non-stratiﬁed diamict in the upstream part of the deposit is interpreted as till by OVS. OCL regard a colluvial origin more likely, based on lack of bimodal distribution of grain sizes, angularity of most of the clasts, and continuity with the dipping stratiﬁed beds. Indirectly, a glacial origin of the diamicts is rendered unlikely by lack of associated abraded pavement, even under bedrock overhangs in the wash bottom, and lack of any other small- or large-scale glacial erosional features in Big Falls wash or the headwalls above it (Figs. 6 and 7). Striations and facets on clasts in the diamict are discussed below. We will next walk up the rough, bouldery channel of Big Falls wash a few hundred meters, to a point below the “Hanging till” of OVS. The exposure itself is high above the channel and the rolling-rock hazard is very high, so we will settle for examination of striated stones at the bottom of the gully that descends from the exposure. The exposure is provided by a slump or washout of an extensive apron of debris that mantles the west slope between the channel and the Monte Cristo cliff (Fig. 7). Interpretation of the diamict here should apply to the whole slope, although OVS seem to refer only to the washout. The exposure at the head of the washout (Fig. 9) shows

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unstratiﬁed diamict consisting of a full spectrum of grain sizes up to about one meter in length, though mostly <0.5 m. Some of the diamict has a silt-sand component while other parts consist of clast-supported gravel. There is a mix of very angular, angular, and subrounded clasts; the ﬁner gravel is mostly angular. Striations occur on some of both angular and subrounded clasts, but there are few, or no, depending on the observer, bullet-shaped clasts, and no striated pavements, bimodal grain-size distribution, or landforms indicative of a glacial environment; the apron of debris appears to be slope colluvium. Clast striations are common in the two exposures described above (Fig. 10), and somewhat polished clast faces that could be interpreted as glacial facets occur in both localities. Striations tend to be short and have widely varying orientations relative to the clast; the work of Atkins (2003) suggests these striations are generated more likely by mass movement than glaciation. However, a minority of the striated clasts do have long striations, most of which are subparallel to the long axes of the clasts. Other things being equal, a glacial interpretation of these is appropriate. Using Occam’s Razor, one could argue that if most of the clast striations appear to be generated by mass movement, then even the glacial-appearing ones probably were generated by mass movement. However, a glacial origin of such striations cannot be ruled out. Some of the polished, apparently faceted clasts in the two localities could be of glacial origin. However, some angular clasts show very fresh, unabraded, approximately planar faces that are obviously fracture surfaces; some transportationabrasion and edge-rounding of these in an alluvial or colluvial setting would produce faces that look like glacial facets. As for the relatively smooth polish that characterizes some of the clast faces, it could presumably be produced by abrasion either in a glacial setting or in an alluvial-bedload or debris ﬂow setting. Problematically, there doesn’t seem to be much available transportation distance to do the job in either kind of environment. If there is a glacial component to the diamicts at the two localities, it is most likely in the form of glacially abraded clasts that have been reworked from older and perhaps higher deposits into colluvial diamicts in which they are now found. But no traces of such older deposits have yet been found. A one-day inspection by Osborn of the higher slopes in the Big Falls drainage, above Big Falls, yielded no evident till, no striated and/or rounded and/or polished clasts like those found below Big Falls, and no evidence of glacial erosion. Hence there is no

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known source of reworked, glacially striated clasts. But the colluvium hypothesis is also incomplete and enigmatic. Striated clasts are found in aprons of apparent colluvium below the Monte Cristo cliff, but so far they do not seem to have come from above the Monte Cristo cliff. If that is indeed the case, it is not clear by what mass movement process the striations were created, how some clasts became polished over a very short distance, and why angular clasts are mixed with subrounded clasts. Hike back to the parking lot. The return to Las Vegas can be either back down Highway 157 to U.S. 95, or via a loop that includes Lee Canyon on the northeast side of Charleston Peak. The following road log is for the Lee Canyon loop; mileages continue from the Mary Jane Falls trail parking lot. 49.5 Junction with Highway 158. Turn left. 50.5 The road passes through shattered rocks associated with a thrust fault, whose upper plate is dominated by Monte Cristo limestone. 53.0 Roadcuts expose Kyle Canyon fanglomerates. In this proximal part of the fan clasts are generally larger than those seen on the drive up Highway 157. 55.8 In the roadcut on the left, the base of the fanglomerates can be seen; gravels bury a subsurface hill of Paleozoic limestone. 56.8 Desert Viewpoint on the right. A short path leads to a view of Las Vegas Valley, which can be considered a topographic expression of the Las Vegas Valley Shear Zone. 58.4 Junction with Highway 156 (Lee Canyon Road). Turn left toward the Lee Canyon ski area. 59.5 Lee Canyon meadow. According to Tingley et al. (2001), there is not another mountain meadow for 150 miles in any direction. The ﬂat silt and sand surface under the meadow may be a consequence of damming of the valley by multiple debris ﬂows on the downstream end. 60.9 Lee Canyon ski area, presently known as the Las Vegas Ski and Snowboard Resort. The ski runs are cut on the north side of the divide that separates Lee Canyon from Kyle Canyon. No one has proposed that Lee Canyon was the site of Pleistocene glaciers. Turn around and descend on the Lee Canyon Road. 63.4 Junction with Highway 158. Continue straight ahead on Highway 156. 77.6 Junction with U.S. 95. Turn right toward Las Vegas. 77.9 To the left is Corn Creek Flat, underlain by extensive sheets of white-to-yellow ﬁne-grained silt and clay deposited from valley-bottom springs in the Pleistocene (Tingley et al., 2001). These deposits were assigned to the Las Vegas Formation by Longwell et al. (1965). From here, return to Las Vegas.

ACKNOWLEDGMENTS Thanks to Steve Rowland, who suggested the ﬁeld trip; Abe Van Luik, Nathan Cronin, and Gennyne McCune, who contributed discussion and ideas; Lewis Owen and Ernie Duebendorfer, who constructively reviewed the original manuscript; and Kristyn “Chiquitita” Adams, who compiled the bibliography. APPENDIX I: OSBORN ET AL. (2008) Problems with a glacial interpretation of landforms and sediments in the Spring Mountains, Nevada Osborn, Jerry, Geoscience Department, University of Calgary, Calgary, Alberta T2N 1N4, Canada; Clark, Douglas, Geology Department, Western Washington University, Bellingham, Washington 98225, USA; Lachniet, Matthew, Geoscience Department, University of Nevada, Las Vegas, NV 89154, USA. Geological Society of America Abstracts with Programs, v. 40 (in press). Whether or not there is a record of late Quaternary glaciation in the Spring Mountains of southern Nevada has been debated since the days of Eliot Blackwelder, and recently Orndorff et al. (2003) (Journal of the Arizona-Nevada Academy of Science, v. 36, p. 37–45) purport to have ended the debate with the discovery of tills east of Charleston Peak, and a claim that bowl-shaped valley heads in the Kyle Canyon drainage are degraded cirques. But tempering the glacial interpretation are the following: 1. Equivocal origin of “cirque” morphology. The bowl shapes and “U-shaped morphology” and “hanging valley” described from the head of Kyle Canyon are common in cliff-and-bench topography developed in horizontal strata in desert regions; there are numerous examples in the Grand Canyon. The hanging valley at Big Falls, and the alleged cirque ﬂoor above, are consequences of a thick resistant layer of limestone in the upper Monte Cristo Formation. There are no smooth headwalls or other evidence of glacial erosion in the Spring Mountains. 2. Lack of glacial depositional landforms. Nowhere in the range are there any of the well-preserved late Wisconsinan moraines so common in the glaciated desert ranges to the north, nor is there evidence of older or partly eroded moraines. 3. Lack of till. Two sediment bodies that were interpreted as till on the basis of striated pebbles are very probably colluvium. One of these, at the mouth of Big Falls wash, consists partly of poorly stratiﬁed diamict, dipping at ~30° NW and containing some openwork gravel. The diamict is likely debris ﬂow deposits. Furthermore, the diamict is located in a stream valley where there is abundant evidence for recent mass wasting. The second body appears to be slope colluvium, partly clast-supported, with a complete spectrum of grain sizes up to one meter. Many stones in the two deposits are striated, and some striations could be glacial in origin, but many are short and deep for their length, and some occur on fresh, non-abraded surfaces and appear to be products of impact rather than grinding. A deﬁnitive interpretation of the diamicts in question remains elusive, considering the mixes of angular and rounded clasts, and different types of striations. But it is at least as likely that rounding and striations were produced during various forms of mass movement, as it is that they are glacial in origin.

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Spenceley, A.P., 2003, Reply to Discussion: Grooves and striations on the Stanthorpe Adamellite: evidence for a possible late Middle–Late Triassic age glaciations: Australian Journal of Earth Sciences, v. 50, p. 1001, doi: 10.1111/j.1440-0952.2003.v50_i6_ain_1.x. Sundell, K.A., 1985, The Castle Rocks Chaos; a gigantic Eocene landslidedebris ﬂow within the southeastern Absaroka Range, Wyoming [Ph.D. thesis]: Santa Barbara, California, University of California, 383 p. Tingley, J.V., Purkey, B.W., Duebendorfer, E.M., Smith, E.I., Price, J.G., and Castor, S.B., 2001, Geologic Tours in the Las Vegas Area: Nevada Bureau of Mines and Geology, University of Nevada, Reno, 140 p. van Hoesen, J.G., and Orndorff, R.L., 2000, Evidence for Pleistocene glaciation in the Spring Mountains, Nevada: Geological Society of America Abstracts with Programs, v. 32, no. 7, p. 16. van Hoesen, J.G., and Orndorff, R.L., 2003, Using GIS to evaluate an enigmatic diamicton in the Spring Mountains, Southern Nevada: The Professional Geographer, v. 55, p. 206–215. van Hoesen, J.G., Orndorff, R.L., and Saines, M., 2000, A preliminary assessment of an ice-contact deposit in the Spring Mountains, Nevada: Geological Society of America Abstracts with Programs, v. 32, no. 6, p. 73. Van Hoesen, J.G., Orndorff, R.L., and Saines, M., 2003, Implications of new evidence for late Quaternary glaciations in the Spring Mountains, Southern Nevada: Journal of the Arizona-Nevada Academy of Science, v. 36, p. 37–45. Van Houten, F.B., 1957, Appraisal of Ridgeway and Gunnison “tillites,” southwestern Colorado: Geological Society of America Bulletin, v. 68, p. 383– 388, doi: 10.1130/0016-7606(1957)68[383:AORAGT]2.0.CO;2. Wentworth, C.K., 1928, Striated cobbles in the Southern States: Geological Society of America Bulletin, v. 39, p. 941–954. Whalley, W.B., and Krinsley, D.H., 1974, A scanning electron microscope study of surface textures of quartz grains from glacial environments: Sedimentology, v. 21, p. 87–105, doi: 10.1111/j.1365-3091.1974.tb01783.x. Winograd, I.J., Coplen, T.B., Jurate M. Landwehr, J.M., Riggs, A.C., Ludwig, K.R., Szabo, B.J., Kolesar, P.T., and Revesz, K.M., 1992, Continuous 500,000-year climate record from vein calcite in Devils Hole, Nevada: Science, v. 258, p. 255–260, doi: 10.1126/science.258.5080.255. Winograd, I.J., Landwehr, J.M., Coplen, T.B., Sharp, W.D., Riggs, A.C., Ludwig, K.R., Kolesar, P.T., 2006, Devils Hole, Nevada, δ18O record extended to the mid-Holocene: Quaternary Research, v. 66, p. 202–212, doi: 10.1016/j.yqres.2006.06.003. Winograd, I.J., Szabo, B.J., Coplen, T.B., and Riggs, A.C., 1988, A 250,000-year climatic record from Great Basin vein calcite; implications for Milankovitch theory: Science, v. 242, p. 1275–1280, doi: 10.1126/science.242.4883.1275. Winterer, E.L., and von der Borch, C.C., 1968, Striated pebbles in a mudﬂow deposit, South Australia: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 5, p. 205–211, doi: 10.1016/0031-0182(68)90114-4. Zamoruev, V.V., 1974, Striations on pebbles and boulders: Lithology and Mineral Resources, v. 9, p. 475–479. Zielinski, G.A., and McCoy, W.D., 1987, Modern snowpack and late Pleistocene equilibrium-line altitudes in the mountains of the Great Basin: Arctic and Alpine Research, v. 21, p. 245–255. MANUSCRIPT ACCEPTED BY THE SOCIETY 16 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 11 2008

Quaternary volcanism in the San Francisco Volcanic Field: Recent basaltic eruptions that profoundly impacted the northern Arizona landscape and disrupted the lives of nearby residents S.L. Hanson* Adrian College, Earth Science Department, 110 S. Madison St., Adrian, Michigan 49221, USA W. Dufﬁeld* Geology Department, Northern Arizona University, Flagstaff, Arizona 86011, USA J. Plescia MS MP3-E104, Applied Physics Laboratory, Johns Hopkins University, 11100 Johns Hopkins Road, Laurel, Maryland 20723-6099, USA

ABSTRACT The San Francisco Volcanic Field, located in northeastern Arizona, is host to over 600 volcanoes. These volcanoes began erupting approximately 6 million years ago in the western portion of the ﬁeld and through time, the locus of activity has migrated eastward. Eruptive products range from basalt to rhyolite, with basalt dominant. Pleistocene vents include Merriam Crater and two associated cinder cones as well as The Sproul, a spatter rampart. One, or several, of these vents produced the Grand Falls ﬂow which spilled over into the Little Colorado River gorge and ﬂowed both up and downstream. Lava ﬁlled the canyon producing a dam and continued to ﬂow ~ 1 km beyond the eastern rim. This changed the course of the river creating the waterfall at Grand Falls. Quaternary volcanism began as a ﬁssure eruption that culminated with the building of Sunset Crater cinder cone. The eruption, which produced a blanket of tephra and two lava ﬂows, was most certainly witnessed by the ancestors of the Pueblo Indians and had a dramatic impact on their lives. The eruption may have caused a shift in population to places such as Wupatki, 30 km to the north, where farming in the arid climate may have been temporarily enhanced by a thin layer of ash that acted as a water-retaining mulch. Melts that produced these dominantly basaltic cinder cones were derived by variable amounts of partial melting of an oceanic island basalt–like mantle source that underwent differing degrees of contamination from the lower crust. Subsequent fractional crystallization of olivine ± clinopyroxene further modiﬁed these melts. Discrete packets of these melts ascended rapidly to produce short-lived volcanic events in the eastern San Francisco Volcanic Field.

*[email protected]; wendell.dufﬁ[email protected] Hanson, S.L., Dufﬁeld, W., and Plescia, J., 2008, Quaternary volcanism in the San Francisco Volcanic Field: Recent basaltic eruptions that profoundly impacted the northern Arizona landscape and disrupted the lives of nearby residents, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 173–186, doi: 10.1130/2008.ﬂd011(08). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Hanson et al. The purpose of this ﬁeld trip is to examine these young cinder cones and their eruptive products in an effort to understand the origin of the eruptions as well as the effects they had on the physiography and native inhabitants of the area. Keywords: Sunset Crater, cinder cone, Grand Falls ﬂow, Arizona

INTRODUCTION The San Francisco Volcanic Field lies on the southern margin of the Colorado Plateau near its boundary with the Basin and Range Province, USA. This ﬁeld, ~5000 km2, extends eastward in an irregular belt from near Williams, Arizona, to the Little Colorado River (~50 km east of Flagstaff) (Fig. 1). The total volume of the San Francisco Volcanic Field is ~500 km3 (Wolfe, 1990) and includes compositions that range from basalt to rhyolite, with basalt dominant. Volcanism began ca. 6 Ma in the western part of the ﬁeld, and the locus of activity has since migrated eastward at an average rate of ~2 cm/yr (Tanaka et al., 1986). Some of the youngest eruptions in the San Francisco Volcanic Field produced Merriam Crater, two nearby unnamed cinder cones, The Sproul spatter rampart (all ca. 20 ka) (Dufﬁeld et al., 2006), and Sunset Crater Volcano (ca. A.D. 1075–1090) (Michael Ort, personal communication, 2007).

The origin of volcanism in the San Francisco Volcanic Field is uncertain. It has been attributed to possible shear heating at the base of the lithosphere (Tanaka et al., 1986), to a possible underlying “ﬁxed” mantle hot spot (Dufﬁeld, 1997), and to possible melting induced by pressure reduction as crustal extension and normal faulting of the Great Basin advanced eastward. This ﬁeld trip will examine these young cinder cones and their eruptive products in an effort to understand the origin of the eruptions as well as the effects they had on the physiography and native inhabitants of the area. ERUPTION HISTORY Merriam Crater, The Sproul, Vent 3036A, and Vent 3036B Merriam Crater, The Sproul, Vent 3036A, and Vent 3036B are located along two northwest en echelon trends and are

Figure 1. Shaded relief map of the San Francisco Volcanic Field. Illumination is from the west. Upper panel shows location of ﬁeld in Arizona. The numerous cinder cones and larger volcanoes are clearly shown.

Quaternary volcanism in the San Francisco Volcanic Field interpreted to have been active at about the same time (Dufﬁeld et al., 2006). Cinders from the Merriam Crater cone, which towers 300 m above the tallest of the other three vents, do not blanket these neighbors, suggesting that Merriam erupted ﬁrst. Lava from one or more of these four vents ﬂowed >10 km to the northeast, where it spilled into the 65-m-deep canyon of the Little Colorado River. Lava ﬁlled the canyon, forming a dam, and advanced ~1 km beyond the eastern rim to feed a tongue-shaped lobe several meters thick (Dufﬁeld et al., 2006) (Fig. 2). The lava dam at Grand Falls is at least 60 m thick; its base is not exposed. Intracanyon lava ﬂowed ~15 km upstream from the dam (Plescia et al., 2001) and ~25 km downstream. The reservoir created behind the new lava dam is now ﬁlled with sediment derived from the upper reaches of the Little Colorado River drainage basin. Today the course of the Little Colorado River, an intermittent stream, meanders across the sediment ﬁll almost directly above the pre-dam position of the canyon and diverts around the margin of the spillover lava lobe before cascading into the original pre-dam channel at Grand Falls (Fig. 2). Post-dam erosion has carved a broad ~7-m-deep river bed along the edge of the overﬂow lobe above the dam. Nearly the entire length of the downstream intra-canyon lava is exposed in the river bed. The upstream equivalent is buried beneath post-dam sediment. Magnetic proﬁles measured transverse to this reach of the river’s course exhibit dipolar anomalies that deﬁne the map position of the now-buried pre-dam canyon (Figs. 3 and 4). An upstream decrease in amplitude of the anomalies is interpreted to reﬂect progressive thinning of the intra-canyon lava (Fig. 5) to a location near proﬁle 11 which presumably

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Figure 2. Geologic sketch map of the Grand Falls lava dam. Note the tongue of lava extending across the old channel line (dashed white line). After the volcanism ended, the Little Colorado formed a new path around the tongue, forming a waterfall back into the incised canyon on the northwest side of the lava tongue.

Figure 3. Location of total magnetic ﬁeld proﬁles to trace the upstream extent of the lava within the old channel.

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Hanson et al. marks the distal end of the buried ﬂow. The overall ﬁrst-order shape of the buried ﬂow is probably a tapering ﬁnger whose bottom surface traces the pre-eruption gradient of the river while the upper surface dips gently downward from the elevation at the top of the lava dam to a somewhat lower elevation beneath the sediment cover at proﬁle 11. A cross section of the canyon-ﬁlling ﬂows is exposed along the northwest margin of the lava ﬂow where it ﬁlls the canyon at Grand Falls. Five emplacement (cooling) units can be recognized in the outcrop. From the base up, the bulk of the dam is formed by a single cooling unit of dense columnar-jointed basalt. Columns are ~50 cm across and form plumose patterns deﬁned by the gently curving axes of “bundles” of columns (Dufﬁeld et al., 2006). Steam, presumably from river water (as fumaroles are absent) percolating through the lava may have produced this irregular pattern of columns by producing an irregular pattern of isotherms. The columnar basalt is overlain by four gently dipping basalt ﬂows, each ~1–2 m thick with a well-deﬁned rubbly vesicular bottoms and tops. These exhibit no differences in physical or chemical weathering of inter-ﬂow surfaces and may simply represent overlapping lobes of a single ﬂow (Dufﬁeld et al., 2006). Paleomagnetic directions measured in samples drilled at the bottom and the top of the exposed section are the same providing more evidence of very little time during growth of the lava dam (Dufﬁeld et al., 2006). Sunset Crater Volcano

Figure 4. Total magnetic ﬁeld proﬁles across the course of the Little Colorado upstream from the dam. Locations shown in Figure 3. Proﬁle 11 indicates the absence of lava; the magnitude of anomaly is an indication of the thickness of the lava at that point.

Figure 5. Magnitude of total ﬁeld anomaly as a function of distance upstream from Grand Falls. Diamonds indicate actual values; dashed line indicates a best ﬁt to those data points.

The initial eruption of Sunset Crater was most certainly witnessed by the ancestors of the today’s Pueblo Indians. It is embedded in Hopi oral history and evidenced by the presence of basalt cobbles with corn impressions in ancient pit houses. These “corn rocks” were made deliberately, possibly by placing corn on a hornito and allowing the spatter to cover it (Elson and Ort, 2003). Volcanic activity began with lava fountaining from a 10-km-long southeast-trending ﬁssure. Several small coeval cones (including Rows of Cones, Gyp Crater, and Vent 512; Fig. 6) were created. The ﬁssure volcanism ceased quickly and eruption became focused at the northwestern end, building the 300-m-high Sunset Crater cinder cone volcano, possibly in several stages. Strombolian eruption there ultimately produced a blanket of tephra covering at least 2000 km2. The eruption was probably visible from a distance of perhaps 30–50 km, with an ash plume visible for hundreds of kilometers (Elson and Ort, 2003). Two lava ﬂows, the Bonito and the Kana’a, effused from the base of the volcano (Holm, 1987; Holm and Moore, 1987) (Fig. 6). Field relations show that, at least in a broad sense, the Bonito lava ﬂow resulted from three pulses of activity (Holm, 1987). Mounds of layered agglutinate, pieces of the early Sunset cone, were rafted on the Bonito Flow for distances of hundreds of meters to the west and northwest (Holm, 1987). The rafted mounds, like deposits at the summit of Sunset Crater volcano, are oxidized red. Subsequent fountaining ﬁlled in the “holes” created by rafting and ultimately produced a nearly axis-symmetric cone.

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Figure 6. Location of very young lava ﬂows in the eastern San Francisco Volcanic Field on a topographic base map. The Bonita and Kana’a ﬂow were erupted from Sunset Crater; several small ﬂows were erupted from Gyp Crater–Row of Cones, and the long ﬂow at the southern end is from Vent 512.

During the waning stages of the eruption, cinders on the rim were oxidized to a red scoria and cemented by silica, gypsum, and iron oxide. This fumarolic activity deposited ﬁve separate incrustations around the central vent of the volcano. These incrustations are composed predominantly of gypsum. Locally, this gypsum is overgrown with opal or a thin coating of sulfur. Numerous accessory minerals, including voltaite, jarosite, magnetite, and hematite, occur throughout (Hanson et al., 2000, 2008). Subsequent to the eruption, agricultural areas near the volcano were left covered with a thick layer of cinder and ash that made it unsuitable for growing crops. Further from the vent, farming may have been enhanced by a thin ash layer that acted as a water retaining mulch (Hooten, et al., 2001; Ort et al., 2002). Thus, the eruption may have caused a shift in the population to places such as Wupatki 30 km to the north, where farming in this arid climate may have been temporarily enhanced.

AGE OF VOLCANISM Because the volcanoes studied on this ﬁeld trip are so young, accurate age determinations have been difﬁcult to obtain. Techniques such as 40Ar/39Ar and K-Ar may yield conﬂicting and inaccurate results because the basalt contains so little potassium and may contain excess argon from mantle or crustal sources (Kelly, 2002). Multiple dating techniques have been applied to try to circumvent these problems. Merriam Crater, The Sproul, Vent 3036A, and Vent 3036B Until recently, the accepted age for the lava dam at Grand Falls was ca. 150 ka, based on the whole-rock K-Ar dating technique (Moore and Wolfe, 1987). However, ﬁeld relations such as the lack of physical and chemical weathering and the

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geomorphic position of the intra-canyon lava ﬂow with respect to current stream base level seemed inconsistent with an age this old. Newly applied techniques have provided a better estimate of the age. The results of infrared stimulated luminescence dating of silty mudstone baked by the lava, 3He cosmogenic exposure ages of samples interpreted to be original ﬂow surfaces, multiple 40 Ar/39Ar step-heating ages, and a paleomagnetic match to the region’s magnetic secular variation curve converge on a much younger age of ca. 20 ka (Dufﬁeld et al., 2006).

(K), rhodonite (Mn), and chromite (Cr). Major elements were analyzed at Michigan State University on a Rigaku SMAX spectrometer. Data were reduced by the fundamental parameter data reduction method (Criss, 1980) using XRFWIN software (Omni Instruments). Trace elements were analyzed at Tulane University using a Seimens 200 X-ray ﬂuorescence spectrometer (XRF) with Cr and Mo anode X-ray tubes.

Sunset Crater

Merriam Crater, The Sproul, Vent 3036A, and Vent 3036B

Entrenched in the Sunset Crater Volcano National Monument and academic literature is the notion that the initial eruption of Sunset Crater volcano began in A.D. 1064–1065 (Smiley, 1958), continued for ~100 yr, and concluded with effusive activity that formed the Bonito lava ﬂow in A.D. 1180. The duration of eruption was based on paleomagnetic data (Champion, 1980; Shoemaker and Champion, 1977). The time of eruption onset is based on the presence of diminished tree ring widths from roof beams in pueblo ruins at Wupatki National Monument, 30 km to the north. However, recently recognized uncertainty of the signiﬁcance of the treering pattern lies in the fact that the provenance of the trees used for dating cannot be determined (Ort et al., 2002), as well as the possibility that factors such as drought, ﬁre, and insect infestation could have produced the observed pattern. Recent additional paleomagnetic studies constrain the onset of eruption to between A.D. 1040 and 1100 (Ort et al., 2002) and do not support a prolonged duration. A pronounced Sr-isotope spike in tree rings within this range suggests that the eruption began between A.D. 1075 and A.D. 1090 (Michael Ort, personal communication, 2007). If the eruption had continued for ~100 yr, one would expect to ﬁnd evidence of weathering and/or erosion within the layered blanket of fallout cinders. Extensive ﬁeld examination has failed to ﬁnd such evidence, leading to the interpretation that eruption was short-lived, consistent with the new and greatly expanded paleomagnetic database noted above. These observations, coupled with the lack of geochemical variation between the cinder eruptions and lava ﬂows, suggest a much shorter eruption duration, perhaps as brief as a month or two to as long as a few years (Ort et al., 2002). Assuming population growth at Wupatki and other nearby pueblos is related to enhanced agricultural conditions subsequent to the eruption of Sunset Crater, the later initial eruption date and shorter duration are more consistent with a population increase beginning in ca. AD 1100.

Basalt from the Grand Falls ﬂow is holocrystalline to slightly hypocrystalline and porphyritic with 5%–10% phenocrysts of euhedral to subhedral olivine. At the toe of the ﬂow, olivine occurs as ≤0.5 mm crystals with a composition of Fo73– , whereas near the dam they are slightly larger (0.5–1.5 mm). 77 Samples range from massive to slightly vesicular, and have <10% by volume vesicles. The groundmass is ﬁne-grained to microcrystalline and is composed predominantly of plagioclase laths (An48–41) with lesser amounts of interstitial and granular clinopyroxene, olivine, and unexolved ilmenite (Hanson, 2006). Near the dam, basalt from the upper ﬂows contains large, subparallel plagioclase laths that are rounded and encased with a reaction rim, suggesting that they are xenocrysts rather than phenocrysts (Dufﬁeld et al., 2006).

PETROGRAPHY

Sunset Crater Basalt from the Bonito and Kana’a ﬂows is vesicular, holocrystalline to hypocrystalline and porphyritic with a ﬁne grained groundmass exhibiting varying degrees of ﬂow texture. Vesicles (≤1 cm) encompass 30%–40% of the volume of the rock near the tops of the ﬂows and decrease in volume and size toward the ﬂow center where they are only several millimeters in maximum diameter and occupy ~20%–30% of the volume of the rock. Phenocryst percentages range from 10% to 20% and are composed of anhedral to subhedral olivine crystals a few millimeters in size and, in a few samples, several millimeter subhedral to anhedral augite. Olivine phenocrysts (Fo87–Fo71) exhibit normal zoning with minor alteration to iddingsite. Plagioclase (An73–An45) exhibits oscillatory and normal core to rim zoning. Aphanitic groundmass material is composed of strongly to weakly parallel plagioclase laths, pale green interstitial augite, and unexolved Fe-Ti oxides. Both of these ﬂows contain xenoliths that are composed primarily of Kaibab Limestone ranging in size from several centimeters to ~10 cm.

ANALYTICAL METHODS GEOCHEMISTRY Minerals were analyzed at the University of New Orleans on an ARL-SEMQ electron microprobe with an accelerating voltage of 15 kV, a beam current of 15 nA, and 60 s counting times. Matrix effects were corrected using a φ(ρZ) correction procedure (Pouchou and Pichoir, 1991). Standards used include labradorite (Na, Al, Si), clinopyroxene (Ca, Fe, Mg, Ti), SrSO4 (Sr), adularia

Merriam Crater, The Sproul, Vent 3036A, and Vent 3036B To date, very little geochemical work has been done on The Sproul, Vent 3036A, Vent 3036B, Merriam Crater, and the Grand Falls ﬂow. Existing analyses include a single analysis from Mer-

Quaternary volcanism in the San Francisco Volcanic Field

Figure 7. Plot of chemical analysis of samples from the eastern San Francisco Volcanic Field for Sunset Crater, Grand Falls, Merriam Crater and the Sproul. Modiﬁed from LeBas et al. (1986) with the basalt ﬁeld subdivided into alkali olivine basalt and tholeiitic basalt.

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Additionally, there is considerable overlap in the composition of the three pulses of lava (Flows 1, 2, and 3) described by Holm (1987). Within these ﬂows, SiO2 content ranges from 46.32 wt% to 47.57 wt% for the oldest ﬂow (Flow 1), from 46.53 wt% to 47.55 wt% for the intermediate (Flow 2), and 47.09 wt% to 47.68 wt% for the youngest ﬂow (Flow 3). Mg numbers for Sunset Crater eruptive products range from 54 to 59, values that are typical for moderately evolved basaltic melts. The Bonito ﬂow, the Kana’a ﬂow, agglutinate mounds, and cinders all exhibit identical trends on a chondrite normalized diagram (Figure 10). Lower LILE and HREEs all suggest a moderate enrichment for the basalt. Radiogenic isotopic ratios for the Bonito and Kana’a ﬂows (87Sr/86Sr = 0.703339 to 0.703413 and εNd = –0.12 to 0.87) suggest the melt source for the Sunset Crater lavas is an oceanic island basalt (OIB)–like parental melt with ~ 30% crustal contamination by lower crustal maﬁc granulites that was subsequently modiﬁed by fractional crystallization of olivine and pyroxene (Blaylock and Smith, 1996). Elevated HFSEs are consistent with a mantle source. DISCUSSION AND CONCLUSIONS

riam Crater cinders (Hooten et al., 2001), a single major element analysis from Merriam Crater and The Sproul (Moore and Wolfe, 1987), and ﬁve from the toe of the ﬂow near Wupatki National Monument (Hanson, 2006). Basalt from the Grand Falls lava ﬂow lies within the tephrite-basanite ﬁeld of the total alkali versus silica (TAS) diagram (Fig. 7), whereas single analyses from The Sproul and Merriam Crater lie in the basalt ﬁeld. These analyses are included on variation diagrams for major elements (Fig. 8) and trace elements (Fig. 9). Mg numbers [100 Mg/(Mg + Fe2+)] are typical for moderately evolved basalts and range from 59 to 60 for the Grand Falls ﬂow (Hanson, 2006). Average chemical compositions were normalized to estimates of their abundance in the primitive mantle and are shown on a spider diagram (Fig. 10). Lower light ion lithophile elements (LILEs) and heavy rare earth elements (HREEs) for samples from the toe of the ﬂow suggest the basalt is moderately enriched. Elevated high ﬁeld strength elements (HFSE) suggest a mantle source. Sunset Crater Whole rock major and trace element compositions are given in Table 1. Basalt from Sunset Crater includes samples from the Bonito ﬂow, the Kana’a ﬂow, and the agglutinate mounds. All have similar compositions and lie in the alkali olivine ﬁeld of a total alkali versus silica diagram (Fig. 7). Variation diagrams for major elements (Figure. 8) and trace elements (Fig. 9) show very little variation within samples from the Bonito and Kana’a lava ﬂows as well as the agglutinate mounds. SiO2 variations are all within analytical error, ranging from 46.32 wt% to 47.80 wt% for the Bonito ﬂow, 46.93 wt% to 47.52 wt% for the Kana’a ﬂow, and 46.94 wt% to 47.79 wt% for the agglutinate mounds (Table 1).

Pleistocene volcanism in the eastern portion of the San Francisco Volcanic Field had profound effects on the landscape and the lives of early native peoples near Flagstaff, Arizona. The Grand Falls ﬂow (ca. 20 ka) spilled into the Little Colorado River gorge, creating a lava dam >60 m thick and continued ~25 km downstream and ~15 km upstream. A lake, now sediment ﬁlled, formed behind the dam. The course of the river was modiﬁed and now, when ﬂowing, spills into the pre-dam gorge, creating a waterfall at Grand Falls. While the source vent for the Grand Falls ﬂow is not known, ﬁeld relations show that it is related to one, or several, of four vents, including Merriam Crater, The Sproul, Vent 3036A, and Vent 3036B. The eruption of Sunset Crater Volcano (ca. A.D. 1075–1090) is intimately entwined with the archaeological history of the area. Stories of the eruption are embedded in Hopi oral history and, even today, the Pueblo Indians consider the volcano sacred. While the eruption may have had spiritual meaning, it also played a part in the migration of these people. A layer of ash, which has long since blown away, may have acted as a water-retaining mulch that temporarily enhanced farming at nearby pueblos. Today, the volcano, lava ﬂows, and tephra look very much like they did 900 years ago. While the processes responsible for melt generation in the San Francisco Volcanic Field are poorly understood, these anorogenic, dominantly basaltic magmas were derived from varying amounts of partial melting of an OIB-like mantle source that underwent variable degrees of contamination from lower crustal material. Subsequent fractional crystallization of olivine ± clinopyroxene further modiﬁed these melts. Blaylock and Smith (1996) used isotopic ratios in an effort to quantify the degree of contamination for the Bonito and Kana’a ﬂows. They suggest the source melt underwent ~30% contamination from lower crustal

Figure 9. Trace element abundances in San Francisco Volcanic Field volcanics plotted against SiO2 wt% for Sunset Crater (squares), Merriam Crater (triangles), The Sproul (diamonds), and the Grand Falls ﬂows (circles).

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Figure 8. Major element abundances plotted as a function of SiO2 wt% for Sunset Crater, Merriam Crater, The Sproul, and the Grand Falls ﬂows.

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Quaternary volcanism in the San Francisco Volcanic Field 17.8 20.9

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Turn left (north) on FR 244A, an unpaved road, and continue to The Sproul. Park.

Stop 1: The Sproul This cigar-shaped vent is one of the four possible sources for the Grand Falls lava dam. The nested spatter ramparts presumably formed as eruption along a ﬁssure waned. The topographic saddle at the southeast end of the ramparts likely marks a breach through which lava escaped to feed a ﬂow headed toward the Little Colorado River. Return to Leupp Road via FR 244A. Figure 10. Trace element diagram normalized to primitive mantle for Sunset Crater, Merriam Crater, and the Grand Falls ﬂow. Data from Sunset Crater cinders are from Hooten et al. (2001). Normalizing factors are from Sun and McDonough (1989).

granulites and was subsequently modiﬁed by fractional crystallization of olivine and pyroxene. High Fo contents for olivine in these basalts represent values that are in equilibrium with mantle peridotite. This, coupled with the presence of xenoliths, xenocrysts, and unexolved ilmenite, is consistent with a rapid ascent to the surface without much evolution. The similar chemistry for the eruptive products within the earlier (Merriam Crater, The Sproul and associated vents) and more recent (Sunset Crater Volcano and the associated ﬁssure) eruptive centers, as well as the lack of weathering horizons between these eruptive products, is consistent with short eruption duration. Thus, it is likely that all of these eruptions were shortlived, resulting from melts that ascended rapidly from the mantle having undergone little evolution. Both the earlier and the more recent events may each represent discrete melting events where pockets of melt rose quickly and were not stored in low-level magma chambers prior to eruption. ROAD LOG Stop locations are shown in Figure 11. Day 1. Grand Falls Lava Dam Depart from the motel in Flagstaff at 8 a.m., and head north out of Flagstaff on U.S. 89. Cumulative mileage Description 0.0 1.3 9.4

Entrance to Elden Trail parking lot, 0.2 mi east of Peaks Ranger Station. Turn east (right) onto the Townsend/Winona Road at the stoplight. Turn left (northeast) onto Leupp Road (Navajo Road 15).

Cumulative mileage Description 24.0 30.3

33.9 37.7 38.9

Turn left (east) on the Leupp Road. Cattle guard at the boundary of the Navajo Reservation. Just a few hundred feet beyond the boundary, turn left onto Navajo Road 70. Navajo Road 6920 comes in from the right, continue on Navajo Road 70. Navajo Road 6910 comes in from the right, continue on Navajo Road 70. Turn left and go ~0.5 mi to the picnic area.

Stop 2: Grand Falls Overlook From this picnic area atop the lava dam, you have an excellent overview of Grand Falls, part of the pre-dam canyon below the falls, and the course of the river upstream of the dam. Lava that overﬂowed the east rim of the canyon advanced ~1 km east of this spot. Intra-canyon lava ﬂowed ~25 km downstream from the falls (obvious outcrops on the walls and bed of the canyon) and ~15 km upstream (lava buried beneath post-dam sediment). Weather and water ﬂow permitting, we will take two short hikes from here. The ﬁrst is down the river bed to the rim of the falls. From the rim of the falls, you have an excellent view of a cross section of most of the lava dam, which includes massive dense columnar basalt capped by four thin lava ﬂows. Bundles of axes of columns exhibit an irregular fanning arrangement. This may be due to the presence of permeating steam during cooling and solidiﬁcation, which produced irregular isothermal surfaces within the lava as the columns were forming. Pillow basalt and hyaloclastite deposits are absent. Perhaps there was very little water in the river when the dam formed, or perhaps these other materials have been eroded away. The falls are developed in Permian Kaibab limestone. Nearby reddish outcrops atop the Kaibab are the Triassic Moenkopi Formation. If the river is not ﬂowing, you can see large potholes where pieces of basalt have eroded the channel. We backtrack through the picnic area to a trail that descends to the base of the falls. Outcrops of the columnar basalt in the river bed indicate that the bottom of the lava dam has yet to be exposed by vertical erosion. Lateral erosion at the base of the falls, however, has moved the east wall of the canyon a few meters east of its pre-dam position.

SC24 46.32 15.66 5.56 0.90 9.96 11.88 8.53 0.17 1.82 0.49 0.32 101.61

SC26 47.57 16.22 3.27 0.81 9.89 11.97 7.99 0.17 1.75 0.47 0.01 100.12

Rb 12.8 14.6 Sr 858 876 Zr 155 158 Ba 545 591 La 42.6 36.5 Ce 81.4 75.8 Y 24.7 24.2 Nb 31.1 31.2 Ta n.d. n.d. V 220 237 Cr 228 168 Co 111 97 Ni 136 116 Pb 4.7 3.8 Th 4 2.5 U n.d. n.d. Cu 78.5 62.9 Zn 112 107 Ga 19.2 18.9 Mo 8.2 6.2 Sn 1.5 2.1 W 287 105 Note: n.d. = not detected.

Oxide SiO2 Al2O3 Na2O K2O CaO FeO* MgO MnO TiO2 P2O5 LOI Sum

12.2 842 151 526 40.4 73.3 23.5 29.8 n.d. 252 240 118 140 n.d. 3.9 n.d. 77.9 115 20.9 5.7 1.8 348

Flow 1 SC39 46.97 15.77 3.21 0.78 9.99 11.83 8.33 0.17 1.79 0.44 0.29 99.58 12 862 151 559 44.6 82.9 24.5 29.9 n.d. 231 209 88.9 120 4.1 3 n.d. 74.5 103 21.5 3.1 1.8 157

SC40 46.82 15.78 3.15 0.76 10.04 11.80 8.30 0.17 1.79 0.44 0.40 99.45 12.2 850 154 548 45.4 88.1 24.1 29.9 n.d. 233 224 89.8 132 3.5 n.d. n.d. 78.7 108 16.7 6.8 n.d. 387

SC49 47.35 16.02 3.18 0.79 10.04 11.83 8.07 0.17 1.77 0.44 0.21 99.87 13 842 152 542 37.7 74.4 24.5 30.9 n.d. 227 224 113 141 3.3 n.d. n.d. 81.8 108 20 4.6 1.9 211

S35A 47.20 15.58 3.09 0.76 10.04 11.81 8.56 0.17 1.80 0.45 0.17 99.63 13.9 880 154 584 41.5 84.9 25.7 30.8 14.9 222 170 76.8 106 4.9 4 n.d. 45.8 97.4 20.4 14.8 1.6 112

Flow 2 SC41 47.55 16.42 3.40 0.85 9.99 11.60 7.78 0.17 1.70 0.46 0.02 99.94 14.3 827 160 554 34.9 85.3 25.4 34.1 n.d. 255 235 115 151 4.5 5.5 5.3 84.9 109 18.6 6.1 1.8 214

SC47 46.53 15.31 3.01 0.85 10.09 12.05 8.75 0.17 1.95 0.47 0.29 99.47 11.7 854 152 545 38.8 77.4 23.7 29.5 n.d. 230 182 79.3 120 4.4 n.d. 4 62.2 109 19.2 5.5 n.d. 163

SC18 47.68 16.11 3.22 0.78 9.90 11.89 8.29 0.17 1.73 0.46 0.07 100.30 12.8 871 156 560 39.9 81.7 24 30.7 n.d. 223 177 111 117 n.d. n.d. n.d. 67.2 114 19.2 5 1.6 165

12.4 864 154 554 33.3 75.1 24.7 31.1 n.d. 225 198 108 112 4.6 n.d. n.d. 81.4 107 20.9 6.7 1.7 115

Flow 3 SC20 SC31 47.63 47.30 16.24 15.92 3.27 3.28 0.80 0.79 9.85 9.96 11.84 11.90 8.11 8.29 0.17 0.17 1.74 1.79 0.46 0.47 0.11 0.08 100.22 99.95 13.4 867 158 574 34.5 65.7 24.2 30.1 n.d. 251 170 113 110 3.8 3.4 n.d. 66.4 107 21.4 6 1.9 176

SC36 47.80 16.21 3.26 0.82 9.82 11.72 7.98 0.16 1.74 0.48 0.07 100.06 11.1 798 139 500 36.5 72.7 22.7 27.1 17.3 202 240 n.d. 111 5 4.6 n.d. 66.8 92.9 16.8 n.d. 1.8 35.4

SC54 47.09 15.93 3.23 0.77 9.98 10.52 8.4 0.17 1.78 0.42 0.19 98.48 12 791 143 503 41.4 87.1 21.6 28.1 n.d. 260 236 n.d. 125 2.7 2.3 n.d. 74.8 99.2 17.5 n.d. 1.6 7.1

Kana’a SC55 46.93 15.81 3.19 0.81 10.09 10.67 8.32 0.17 1.83 0.44 0.43 98.69

TABLE 1. WHOLE-ROCK MAJOR AND TRACE ELEMENT DATA FOR SUNSET CRATER VOLCANO

14 864 153 523 34.6 82.4 25.3 30.7 n.d. 209 220 n.d. 126 6.1 4.8 n.d. 98.2 95.7 18.6 n.d. 2.6 n.d.

SC56 47.03 15.65 3.22 0.78 10.08 10.94 8.43 0.17 1.81 0.43 0.15 98.69 12.7 873 152 517 35.6 66.6 24 28.4 n.d. 221 170 88.5 118 4.5 4.5 n.d. 67.8 106 20.4 6.6 1.5 105

SC23 47.41 16.33 3.34 0.77 9.79 12.03 8.12 0.17 1.78 0.42 0.01 100.17

13.1 856 157 548 47.2 80.3 24.8 30.6 n.d. 224 197 123 115 3.4 4.4 n.d. 66 113 20.5 6.3 1.2 131

SC44 46.94 16.13 3.05 0.82 9.91 11.97 7.95 0.17 1.80 0.45 0.76 99.95 14.9 889 161 618 48.7 84.5 24 31.2 n.d. 231 153 67.8 100 5.8 4.3 5.4 68.5 108 20.5 5.6 1.1 146

Agglutinate SC37 47.79 16.10 3.25 0.80 9.87 11.86 8.14 0.17 1.78 0.46 0.01 100.23

182 Hanson et al.

Quaternary volcanism in the San Francisco Volcanic Field

Figure 11. Map showing locations of ﬁeld trip locations.

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Hanson et al.

Hike back to the picnic area. Time and weather permitting, we will drive ~6 mi upstream from Grand Falls to visit an older lava dam. Stop 3: Overlook View into an Abandoned Reach of the River and the Downstream End of the Lava Dam that Caused Abandonment The Little Colorado River has been dammed by basalt lavas of the San Francisco Volcanic Field at several locations. Formation of two such dams at ~30 km and 80 km downstream of Grand Falls occurred ca. 2.4 Ma and 0.5 Ma, respectively, on the basis of K-Ar whole-rock ages reported by Damon et al. (1974). The dam seen here is ca. 175 ka (William McIntosh, 2007, personal communication) based on 40Ar/ 39Ar step-heating age. It is in a part of the river canyon eroded across the crest of a gentle anticline in the Moenkopi and older bedrock. We’re standing on Kaibab limestone of that structure. A large lake likely formed behind the lava dam and the anticline, until water overtopped the crest of the fold and eroded a new three-mile segment of canyon that connects with the temporarily abandoned downstream channel of the river. By looking downstream, to the left, we can see where the new canyon and this abandoned reach join. Drive back to Flagstaff. Day 2. Sunset Crater Volcano Depart from the motel in Flagstaff at 8 a.m. Drive north on U.S. 89 for ~15 mi to the intersection with FR 545 (the Sunset Crater–Wupatki loop road). Set the trip odometer to 0.0 at this intersection. Cumulative mileage Description 0.0 1.1

Intersection of U.S. 89 and FR 545. Bonito Pullout—San Francisco Peaks.

Stop 1: View of San Francisco Peaks This stop provides an excellent view of the San Francisco Peaks to the west and Sunset Crater Volcano to the east. Both of these mountains are sacred to the local native people. The Hopi consider the San Francisco Peaks home to the spirit messengers, the Katsina, and to the Navajo it is one of the four sacred mountains that border their spiritual homeland. This volcano, the only composite volcano in the San Francisco Volcanic Field, began erupting ca. 2.8 Ma and concluded with the emplacement of Sugarloaf dome (ca. 91 ± 2 ka) based on 40Ar/39Ar step heating (W. McIntosh, 2007, personal commun.). The high peak on the north side of the mountain is Mount Humphreys, the highest point in Arizona (3850 m; 12,633 ft.). However, a large portion of this mountain may have been lost during either an explosive eruption or landslide activity. The original height of the volcano may have been as great as 4900 m (16,000 ft.) (Dufﬁeld, 1997). Sunset Crater Volcano, having erupted in ca. A.D. 1075– 1090, was most certainly witnessed by the ancestors of today’s

Pueblo tribes. Even today, the Hopi consider Sunset Crater to be the home of the Hopi Kana’a Katsina and have the story of the eruption embedded in their oral history. Continue north on FR 545. Cumulative mileage Description 1.9

Park in the Sunset Crater Visitor Center parking lot.

Stop 2: Visitor Center The Visitor Center has excellent exhibits on both the science and the cultural impacts of the eruption of Sunset Crater eruption, so we will take a little time to peruse the exhibits and purchase souvenirs here. Continue north on FR 545. Cumulative mileage Description 3.2

Turn left into the Lenox Crater Trailhead and Lava Flow Vista parking area.

Stop 3: Lenox Crater Trailhead and Lava Flow Vista At this pullout, you are standing on the edge of the Bonito Lava ﬂow. This ﬂow effused from the northwestern base of the volcano and ponded in an intercone basin covering an area of 4.63 km2 to a depth ranging from ~2 m locally along the margin to perhaps >30 m in the center (Moore, 1974; Holm, 1987). From here we will walk the short trail down onto the Bonito Lava ﬂow to the end of the trail where a large slab of lava was turned up on its side. Visible behind the ﬂow are two dome volcanoes, O’Leary Peak and Robinson Dome, which are composed primarily of dacite. These volcanoes were active from approximately 250 to 170 ka (Moore and Wolfe, 1987). Continue north on FR 545. Cumulative mileage Description 3.4

Turn right into the Lava Flow Trail parking area.

Stop 4a: Lava Flow Trail We will walk the 1 mile loop trail through the Bonito Lava ﬂow. From the trail you can see each of the three ﬂow units described by Holm (1987). Holm (1987) designated these units as Flows 1, 2, and 3 with Flow 1 the oldest and Flow 3 the youngest. Each of the three distinct extrusive stages is characterized by a particular relative height of the top of the ﬂow and the degree to which it is covered by cinders from Sunset Crater. The stratigraphy is inverted as the lowest elevation ﬂows have virtually no cinder cover, whereas those at higher elevation have considerable cinder cover. Stop 4b: Squeeze-Up Squeeze-ups are prevalent, especially along the margin of the Bonito Lava ﬂow.

Quaternary volcanism in the San Francisco Volcanic Field Stop 4c: Ice Cave At the far end of the Lava Flow Trail is a subterranean feature that has been interpreted as either a lava tube or a dilational ﬁssure that did not reach the surface. This cave has irregular walls of blocky slumped material and continues back for several hundred feet. The entry is gated because of a partial collapse in the early 1990s. Stop 4d: Hornito and Spatter Several hornitos, rootless spatter cones, are present on the Bonito and Kana’a ﬂows. The hornito next to the Lava Flow Trail is the largest on the Bonito Lava ﬂow. It was on a spatter rampart like this that the ancestors of the Pueblo Indians may have placed corn to create the corn rocks found in nearby pit houses (Elson et al., 2002). Stop 4e: Agglutinate Mounds The red mounds across the road are agglutinate material that was rafted on the Bonito Lava ﬂow for several hundreds of meters to the west and northwest from the early summit of the volcano. Continue north on FR 545. Cumulative mileage Description 5.4 5.5

Turn east into the Cinder Hills overlook and drive ~0.1 miles to the overlook. Park at the Cinder Hills overlook.

Stop 5: Cinder Hills Overlook The Cinder Hills are composed of a number of older cinder cones that erupted hundreds of thousands of years ago. Strombolian eruptions from Sunset Crater produced a layer of tephra that initially covered an area of ~2300 km2 to >1 cm depth (Elson et al., 2002). This tephra was blown by the prevailing winds in a predominantly easterly direction, and thus blanketed the older cinder cones. Although much of the ﬁner ash has long since blown away, its former presence is indicated by human history. A thriving community began to ﬂourish 30 km to the north at Wupatki immediately following the eruption. Archaeologists have suggested that this may be because a thin layer of ash could have acted as a water-retaining mulch that allowed for farming in an area that was previously too dry (Hooten et al., 2001). Extending to the east-southeast from Sunset Crater (to your south) lies a thin, linear trace of red cinders and agglutinate. This represents the trace of the ﬁssure eruption. Return 0.1 miles to FR 545 and continue north on FR 545. Cumulative mileage Description 8.1

Park at the pullout on the east side of the road.

Stop 7: Kana’a Flow This ﬂow, described in detail by Holm and Moore (1987), was extruded from the eastern base of the volcano and ﬂowed

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in a northeasterly direction following Kana’a wash. It ﬂowed for ~8 km down the wash, yet is only ~350 m wide. We will walk a short distance (<100 m) off the trail to examine the Kana’a ﬂow. End of road log. ACKNOWLEDGMENTS We would like to thank the National Park Service, Adrian College, and the GSA GeoCorps America™ program for ﬁnancial support of this research. Additionally, we would like to thank Carol Mann and the John F. Mann Memorial Fund and Larry Wooden from the Shell Exploration and Production Company for supporting the GeoCorps America program and making the ﬁrst author’s ﬁfth summer at Sunset Crater Volcano National Monument possible. We are grateful to Tom Vogel at Michigan State University for the laser ablation–inductively coupled plasma–mass spectrometer analyses, to Pierre Burnside at Tulane University for the X-ray ﬂuorescence spectrometer analyses, and to Al Falster at the University of New Orleans for performing the electron probe microanalyses. We also owe thanks to Kelly O’Connell for help with the ﬁeld work at Grand Falls. REFERENCES CITED Blaylock, J., and Smith, E., 1996, Geochemical investigation of Sunset Crater, Arizona: Complex history of a low-volume magmatic system: Geological Society of America Abstracts with Programs, v. 28, no. 7, p. A-162. Champion, D.E., 1980, Holocene geomagnetic secular variation in the western United States: Implications for the global geomagnetic ﬁeld: U.S. Geological Survey Open-File Report 80-824, 314 p. Criss, J.W., 1980, Fundamental parameters calculations on a laboratory microcomputer: Advances in X-Ray Analysis, v. 23, p. 93–97. Damon, P.E., Shaﬁqullah, M., and Leventhal, J.S., 1974, K-Ar chronology for the San Francisco volcanic ﬁeld and rate of erosion of the Little Colorado River, in Karlstrom, T.N.V., et al., eds., Geology of Northern Arizona, with notes on archaeology and paleoclimate; Part 1, Regional studies: Northern Arizona University, p. 221–235. Dufﬁeld, W.A., 1997, Volcanoes of northern Arizona: Arizona, Grand Canyon Association, 68 p. Dufﬁeld, W.A., Riggs, N., Kaufman, D., Champion, D., Fenton, C., Forman, S., McIntosh, W., Hereford, R., Plescia, J., and Ort, M., 2006, Multiple constraints on the age of a Pleistocene lava dam across the Little Colorado River at Grand Falls, Arizona: Geological Society of America Bulletin, v. 118, p. 421–429, doi: 10.1130/B25814.1. Elson, M.D., and Ort, M.H., 2003, Collaborative research at Sunset Crater Volcano, in Elson, M.D., ed.: Center for Desert Archaeology, Archaeology Southwest, v. 17, p. 4–6. Elson, M.D., Ort, M.H., Hesse, S.J., and Dufﬁeld, W.A., 2002, Lava corn and ritual in the northern southwest: American Antiquity, v. 67, p. 119–135, doi: 10.2307/2694881. Hanson, S.L., 2006, Characterization and correlation of lava ﬂows in Wupatki National Monument, northern Arizona: Western National Parks Association Research Report no. 06-11, 12 p. Hanson, S.L., Falster, A.U., and Simmons, W.B., 2000, Mineralogy of fumarole deposits from Sunset Crater Volcano, northern Arizona: Rochester, New York, Rochester Mineralogical Symposium Abstracts, p. 9. Hanson, S.L., Falster, A.U., and Simmons, W.B., 2008, Mineralogy of fumarole deposits at Sunset Crater Volcano National Monument, Northern Arizona: Rocks and Minerals (in press). Holm, R.F., 1987, Signiﬁcance of agglutinate mounds on lava ﬂows associated with monogenetic cones: An example from Sunset Crater, northern Arizona: Geological Society of America Bulletin, v. 99, p. 319–324, doi: 10.1130/0016-7606(1987)99<319:SOAMOL>2.0.CO;2.

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Holm, R.F., and Moore, R.B., 1987, Holocene scoria cone and lava ﬂows at Sunset Crater, northern Arizona: Geological Society of America Centennial Field Guide—Rocky Mountain Section, p. 393–397. Hooten, J.A., Ort, M.H., and Elson, M.D., 2001, Origin of cinders in Wupatki National Monument: Desert Archaeology, Inc., Technical Report No. 2001-12, 20 p. Kelley, S. P., 2002. Excess argon in K-Ar and Ar-Ar geochronology. Chemical Geology, v. 188, p. 1–22. LeBas, M.J., LeMaitre, R.W., and Streckeisen, A.L., 1986, A Chemical Classiﬁcation of Volcanic Rocks Based on the Total Alkali-Silica Diagram: Journal of Petrology, v. 27, p. 745–750. Moore, R.B., 1974. Geology, petrology and geochemistry of the eastern San Francisco Volcanic Field, Arizona [Ph.D. thesis]: Albuquerque, New Mexico, University of New Mexico, 360 p. Moore, R.B., and Wolfe, E.W., 1987, Geologic map of the east part of the San Francisco volcanic ﬁeld, north-central Arizona: U.S. Geological Survey Map MF-1960, scale 1:50,000. Ort, M.H., Elson, M.D., and Champion, D.E., 2002, A paleomagnetic dating study of Sunset Crater Volcano: Desert Archaeology, Inc., Technical Report No. 2002-16, 16 p. Plescia, J.B., Dufﬁeld, W.A., McIntosh, W.C., Forman, S.L., and Champion, D.E., 2001, Timing and extent of Grand Falls lava, Little Colorado River: Sixth Biennial Conference of Research on the Colorado Plateau, p. 46–47.

Pouchou, J.-L., and Pichoir, F., 1991, Quantitative analysis of homogenous or stratiﬁed microvolumes applying the model “PAP,” in Heinrich, K.F.J., and Newbury, D.E., eds., Electron Probe Quantitation: New York, Plenum, p. 31–75. Shoemaker, E.M., and Champion, D.E., 1977, Eruption history of Sunset Crater, Arizona: Investigator’s Annual Report: Manuscript on ﬁle, Flagstaff Area National Monuments Headquarters, Wupatki, Sunset Crater Volcano, and Walnut Canyon National Monuments, Flagstaff, Arizona. 5 p. Smiley, T.L., 1958, The geology and dating of Sunset Crater, Flagstaff, Arizona, in Anderson, R.Y., and Harshbarger, J.W., eds., Guidebook of the Black Mesa Basin, northeastern Arizona: Socorro, New Mexico, New Mexico Geological Society, p. 186–190. Sun, S., and McDonough, W.F., 1989, Chemical and isotopic systematic of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the Ocean Basins: Boston, Blackwell Scientiﬁc, p. 313–345. Tanaka, K.L., Shoemaker, E.W., Ulrich, G.E., and Wolfe, E.W., 1986, Migration of volcanism in the San Francisco volcanic ﬁeld, Arizona: Geological Society of America Bulletin, v. 97, p. 129–141, doi: 10.1130/00167606(1986)97<129:MOVITS>2.0.CO;2. Wolfe, E.W., 1990, San Francisco, Arizona, in Wood, C.A., and Kienle, J., eds., Volcanoes of North America: Cambridge, Cambridge University Press, p. 278–280. MANUSCRIPT ACCEPTED BY THE SOCIETY 17 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 11 2008

The Spirit Mountain batholith and Secret Pass Canyon volcanic center: A cross-sectional view of the magmatic architecture of the uppermost crust of an extensional terrain, Colorado River, Nevada-Arizona Nicholas P. Lang Department of Earth and Planetary Sciences, University of Tennessee, Knoxville, Tennessee 37996, USA B.J. Walker Department of Geosciences, Oregon State University, Corvallis, Oregon 97331, USA Lily L. Claiborne Calvin F. Miller Department of Earth and Environmental Sciences, Vanderbilt University, Nashville, Tennessee 37235, USA Richard W. Hazlett Geology Department, Pomona College, Claremont, California 91711, USA Matthew T. Heizler New Mexico Bureau of Geology and Mineral Resources, New Mexico Tech, Socorro, New Mexico 87801, USA

ABSTRACT Extreme extension along the Colorado River has exposed the shallow to midcrustal Spirit Mountain batholith and the roots of the roughly coeval Secret Pass Canyon volcanic center. Examination of the Spirit Mountain batholith reveals evidence for multiple replenishment and rejuvenation over a two million year period (ca. 17.5– 15.3 Ma), with extensive coarse cumulate granites and leucogranite (high-silica rhyolites) sheets, maﬁc-felsic mingling and mixing, and a major dike swarm. The roots of the possibly related Secret Pass Canyon volcanic center comprise a large, very shallow, composite laccolith and smaller dikes, sills, and a volcanic neck. The volcanic sequence was emplaced within about a one million year period (ca. 18.5–17.3 Ma) and includes volcanogenic sediments, ignimbrites, domes, and block-and-ash ﬂow deposits. An appended road log serves as a geologic guide to this magmatic region. Keywords: Colorado River extensional corridor, magmatic plumbing, batholith, volcanic center, Basin and Range

Lang, N.P., Walker, B.J., Claiborne, L.E., Miller, C.F., Hazlett, R.W., and Heizler, M.T., 2008, The Spirit Mountain batholith and Secret Pass Canyon volcanic center: A cross-sectional view of the magmatic architecture of the uppermost crust of an extensional terrain, Colorado River, Nevada-Arizona, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 187–214, doi: 10.1130/2008.ﬂd011(09). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION Nevada

The northern Colorado River extensional corridor is a 70to 100-km-wide region of moderately to highly extended crust along the eastern margin of the Basin and Range province in southern Nevada and northwestern Arizona (Fig. 1) (Howard and John, 1987; Faulds et al., 1990, 2001). The corridor is dominated by moderately to steeply tilted fault blocks that contain thick Miocene volcanic sequences. Dissection by the Colorado River and its tributaries has generated spectacular exposures of extensional structures and Miocene volcanic and plutonic complexes (e.g., Anderson, 1971). Some of the larger, more steeply tilted fault blocks expose thick crustal sections, including parts of Miocene volcanic ediﬁces, hypabyssal dike swarms, and large midto upper-crustal plutonic complexes. Consequently, the plumbing systems of some Miocene magmatic systems can be analyzed down to mid-crustal levels (e.g., Faulds et al., 1995; Bachl et al., 2001; Walker et al., 2007). The stratigraphy of the northern Colorado River extensional corridor is characterized by thick sections (generally >3 km) of Miocene volcanic and sedimentary strata that rest directly on Proterozoic and Late Cretaceous metamorphic and plutonic rock (Anderson, 1971, 1977, 1978; Faulds, 1995, 1996; Faulds et al., 1990, 1995, 2001; Sherrod and Nielson, 1993). Miocene volcanic sections include felsic to maﬁc lavas, volcanic breccia, and several regionally extensive ash-ﬂow tuffs. Magmatism and large-magnitude ~E-W extension swept northward through the corridor in early to middle Miocene time (Faulds et al., 2001). Magmatism began ca. 20 Ma in the south (southern Black Mountains) and ca. 13 Ma in the north (western Lake Mead area), generally preceding the onset of major eastwest extension by 1–4 m.y. Voluminous intermediate lava characterized the early phase of volcanism, which was accompanied in many areas by mild north-south extension, as evidenced by abundant coeval ~east-striking dike swarms and normal faults. The onset of major ~east-west extension occurred ca. 16 Ma in

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GEOLOGICAL SETTING: THE NORTHERN COLORADO RIVER EXTENSIONAL CORRIDOR

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The purpose of this ﬁeld trip is to examine cross-sectional views of a Miocene batholithic complex and roughly coeval eruptive center in southeastern Nevada and adjacent northwestern Arizona, respectively. We will emphasize the storage transport, fractionation, and eruption processes at a range of depths exposed by extension-related tilting. The shallow to mid-crustal Spirit Mountain batholith reveals evidence for multiple replenishment and rejuvenation over a 2 m.y. period. Across the Colorado River, the roots of the possibly related Secret Pass Canyon volcanic center comprise a large, very shallow, composite laccolith and smaller dikes, sills, and a volcanic neck, with intriguing ﬁeld relationships with the volcanic sequence that it intrudes and underlies. The sequence includes volcanogenic sediments, ignimbrites, domes, and block-and-ash ﬂow deposits.

SPCVC Kingman

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Figure 1. Colorado River Extensional Corridor in the California-Nevada-Arizona tri-state region. Day 1 of this trip will be spent examining the Spirit Mountain batholith (SMB) in the Newberry Mountains, and Day 2 will be in the Black Mountains examining the Secret Pass Canyon volcanic center (SPCVC); dashed black boxes show the locations of Figure 9. LMSZ—Lake Mead shear zone; LVVSZ—Las Vegas Valley shear zone.

the southern to central parts of the region and progressed northward to the western Lake Mead area by ca. 13 Ma. During peak extension ca. 15–16 Ma, the early intermediate volcanism generally gave way to bimodal compositions (Faulds et al., 1995). In most areas, felsic volcanism ended abruptly during or shortly after peak extension. However, widespread maﬁc magmatism continued through the waning stages of extension (Gans and Bohrson, 1998; Feuerbach et al., 1999; Faulds et al., 2001). After extension ceased ca. 12–8 Ma, from south to north across the region, magmatism dwindled to localized eruptions of tholeiitic and alkalic basalt, which are as young as ca. 4 Ma in the northern part of the corridor near Lake Mead (Smith et al., 1990; Feuerbach et al., 1993). Exposed plutons range between ca. 18 and ca. 13 Ma and were emplaced 2–4 m.y. after volcanism began, generally near to or spanning onset of rapid extension (Miller et al., 2005). The Spirit Mountain batholith, which we will investigate on this trip, is the largest intrusive complex and spanned most of the ﬁrst half of this interval (ca. 17.4–15.3 Ma). As with other major plutons, the Spirit Mountain batholith documents both maﬁc and felsic input into the upper crust and, though dominated by granite, contains rocks with a very wide range of compositions. Near-surface

Spirit Mountain batholith and Secret Pass Canyon volcanic center hypabyssal intrusions and associated extrusive rocks across the river at Secret Pass Canyon volcanic center were emplaced at the onset of major plutonism, coeval with early stages of volcanism (ca. 18.5–17.4 Ma). Here, compositions are almost all intermediate to felsic (silicic andesite to high-Si rhyolite), but there are hints that maﬁc magma was also involved in the system. SPIRIT MOUNTAIN BATHOLITH AND SECRET PASS CANYON VOLCANIC CENTER: VIEWS OF MAGMATISM IN THE UPPER CRUST The connections between volcanic and plutonic systems are not well understood, and some of the best insights available are likely to be in shallow intrusive complexes. The Spirit Mountain batholith and Secret Pass Canyon volcanic center provide near– cross-sectional exposures of varying levels of subvolcanic to shallow batholithic complexes. Although presently the two areas lie on opposite sides of the Colorado River, with volcanics and very shallow intrusives exposed in Secret Pass Canyon volcanic center east of the river and deeper plutons of the Spirit Mountain batholith exposed on the west side, inferred large-scale eastward transport of allochthonous upper crust on the regional detachment system (e.g., Faulds et al., 2000, 2001) permits rough proximity of the two at the time when they were magmatically active. In this view, the Secret Pass Canyon volcanic center was in the uppermost crust whereas the Spirit Mountain batholith would represent the lower portion of the upper crust and upper mid-crust during the period of intense magmatic activity. The view of two levels of exposure in a very magmatically active area can provide insight to the variety of mechanisms for magma transport and storage in the shallow crust. The Spirit Mountain batholith represents gradual accumulation of magma in a “patchwork” batholith over 2 m.y. (from ca. 17.4 to 15.3 Ma) by repeated intrusion. Field evidence suggests intruding magma often ponded as sill-like structures within weak zones of the pluton (presumably melt-rich zones) merged to form thick sequences of granite. The onset of voluminous activity commenced at the Secret Pass Canyon volcanic center around 18.5 Ma, when voluminous trachydacite, which we interpret as an extremely shallow, ﬂat-topped sill-laccolith complex, was emplaced into a sequence of tuffs and volcanogenic sediments that may only be very slightly older. Subsequently, over a short period between ca. 17.3 and 17.7 Ma, at least one silicic ignimbrite and thick dome-derived rhyolite ﬂows and breccias were deposited and then intruded by andesitic to dacitic dikes, sills, and volcanic necks. Despite the paleotectonic regime and the overlap in some ages, no evidence has surfaced to deﬁnitively connect the two magmatic systems. The Spirit Mountain Batholith intrudes Proterozoic gneisses, and although dikes that extend upward from the roof of the batholith have been identiﬁed, nowhere do either batholiths or dikes intrude volcanic rocks. The onset of Secret Pass Canyon volcanic center activity appears to have predated that of the Spirit Mountain batholith by a million years, followed by copious eruption of rhyolite at the Secret Pass Canyon

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volcanic center that coincide with the early stages of the Spirit Mountain batholith, permitting a possible linkage that we plan to explore further. The Spirit Mountain Batholith For years, Spirit Mountain has been regarded as something of a mystery. Native American legends called it Avikwa’ame and hail it as the birthplace of local Mojave tribes. It is the place where the mighty god Mastamho is said to have piled up earth in a mourning ceremony for his recently slain father. Mastamho then brought the Mojave people to this land and built his house at the summit of Avikwa’ame (Spirit Mountain), where he established himself as shaman-lord. Native American shaman aspirants are said to have gone on vision quests at the foot of Spirit Mountain in an attempt to revisit the creation events. Hints of this myth have lived into the present day—local residents still recall stories of a shaman who lives atop Spirit Mountain. Thousands of petroglyphs adorn the rocks surrounding Spirit Mountain, providing evidence that this was a well-visited, if not indeed sacred, landscape. Geologically speaking, Spirit Mountain and the surrounding area remained unstudied until the mid-1960s. Volborth (1973) ﬁrst mapped and described in reconnaissance the “Spirit Mountain block” in the Newberry Mountains at the southern tip of Nevada. Subsequent studies (e.g., Hopson et al., 1994; Howard et al., 1994; Faulds et al., 1992; Haapala et al., 2005) established the Spirit Mountain pluton and the adjacent Mirage pluton as tilted granite complexes. Based on chemistry, geochronology, and ﬁeld relations, these two plutons have since been considered collectively as the Spirit Mountain batholith. Preliminary geochronological studies of the Spirit Mountain and Mirage plutons showed them to be middle Miocene (Howard et al., 1996; Ramo et al., 1999). The Spirit Mountain batholith intruded three extant units: a 1.7 Ga gneiss complex, a 1.4 Ga megacrystic granite, and the Late Cretaceous White Rock Wash Pluton. After emplacement, as a consequence of extension in the northern Colorado River extensional corridor, it was tilted 40–50° westward, as indicated by paleomagnetic data (Faulds et al., 1992) and west-east progression from miarolitic leucogranite to coarse-grained, foliated quartz monzonite (Hopson et al., 1994). This tilting affords a cross section of the batholith in map view, with a westward paleo-up direction (Fig. 2). Textural and Compositional Gradation in the Spirit Mountain Granite By far the largest unit of the Spirit Mountain batholith is the Spirit Mountain granite. It constitutes a sequence that ranges from west to east (top to bottom), for the most part gradationally, from high-silica leucogranite to granite to foliated quartz monzonite. The gradation in this massive unit is seen texturally and compositionally, with few, if any, gaps. Granitoid samples with <~70 wt% SiO2 show textural evidence for accumulation of feldspars, biotite, and accessory minerals, and they are enriched in elements

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Fine Grained Granite w/ local blocks/screen of SM granite “Mirage Pluton”: fine-med grained granite, undivided

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BC101z Zircon sample location x (results shown in Fig. 4)

Figure 2. Geologic map of the Spirit Mountain (SM) batholith. Up direction prior to tilting indicated by arrow labeled “Paleo-up.” Highway 163 shown for reference. The Newberry Mountains dike swarm cuts through the central portion of the batholith, but is absent to avoid visual complication. Simpliﬁed from Walker et al. (2007). Data obtained from zircon samples labeled on the map are shown in Figure 4.

that would be concentrated by accumulation (Fig. 3): Ca, Al, Ba, Sr, and Eu (feldspar accumulation), Fe and Mg (biotite and oxide mineral accumulation), and Ti, P, Zr, and light rare earth elements (REE) (accessory sphene, apatite, zircon, allanite, and chevkinite accumulation). In contrast, the most silicic rocks are low in all of these elements and extremely depleted in Sr and Ba (to <10 ppm and <20 ppm, respectively) and Eu. Very low Zr-Hf (17–30 ppm) in these highly silicic rocks is consistent with fractionation of zircon (Claiborne et al., 2006). Geochemically, the Spirit Mountain granite end members—high silica leucogranite and quartz monzonite—are dead ringers for a cumulate compaction/melt segregation model as proposed by Bachmann and Bergantz (2004). The gradational Spirit Mountain granite is cut by several smaller intrusive units. Reconnaissance investigations indicate that the Mirage granite is largely ﬁne- to medium-grained in part porphyritic granite with a minor maﬁc component (Walker et al., 2007; cf. Haapala et al., 2005; Howard et al., 1996). Possibly

related to the Mirage rocks are a ﬁne-grained granite and a diorite that cut the Spirit Mountain granite as a series of (initially) subhorizontal sheets and the Newberry dike swarm, which intruded vertically. The ﬁne-grained granite and the felsic Newberry dikes have almost indistinguishable elemental chemistry, and much of Mirage granite is also similar, suggesting that later felsic magma input into the Spirit Mountain batholith was highly uniform. These units may, in fact, represent a common input magma for the Spirit Mountain batholith system, earlier injections of which differentiated to form the observed gradational sequences in the coarser granitoids. Zircon Geochronology and Geochemistry Fifteen zircon samples from the Spirit Mountain batholith were dated using U-Pb methods at the sensitive high-resolution microprobe–reverse geometry (SHRIMP-RG) lab at Stanford University; we present results for four of those samples in

Spirit Mountain batholith and Secret Pass Canyon volcanic center SM granite

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Figure 3. Harker plots of chemical trends of various units within the Spirit Mountain (SM) batholith.

Figure 4. The zircon data demonstrate that the Spirit Mountain batholith had an assembly time of ~2 m.y. (ca. 17.4–15.3 Ma). Furthermore, most samples (including all Spirit Mountain granite samples) yielded complex age data, showing multiple populations and ca. 2 Ma spectra (Fig. 4). We dated nine samples from the Spirit Mountain granite, which all yielded at least two zircon age populations. In most samples, the populations that make up the minor peaks consist of older grains (or areas within grains), which we interpret to be recycled crystals, or antecrysts.

The common presence of antecrysts indicates that an appreciable fraction of zircon, and presumably other phases as well, were recycled into or affected by subsequent injections. In some cases, smaller, young peaks in an age spectra may reﬂect (minor?) episodes of mush reactivation. Elemental zoning within the zircons, which documents major ﬂuctuations in temperature and host melt composition, strongly supports this complex history for individual zircons and zircon populations (Hf, U, Th, REE, and Ti-in-zircon thermometry; Claiborne et al., 2006; Watson and

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Figure 4. Probability density plots for four Spirit Mountain batholith zircon samples. Vertical lines behind the histogram indicate age populations, each of which is labeled with a date that was established by UNMIX (after Sambridge and Compston, 1994) in Isoplot. Abbreviations SWz, BGz, LGz, and BC101z refer to sample numbers; FGG—ﬁnegrained granite; SMG—SM granite; SML—SM leucogranite; SMQM—Spirit Mountain quartz

Harrison, 2005). These zircon data indicate that the gradational Spirit Mountain granite does not represent a single stage, crystal mush cumulate–melt segregation event. Under close scrutiny, ﬁeld relations also contradict this simple scenario and are instead consistent with a multistage intrusive history. The Gradational Spirit Mountain Granite Widespread crystal recycling suggests an explanation for the rarity of sharp contacts in a unit whose crystallization age spans 2 m.y. Intrusion of fresh magma could remobilize an area within a stagnant, crystal-rich mush pile. The resulting physical interaction might obscure evidence of the injection and erase earlier contacts as well. We envision a large pile of crystal-rich mush that compacted under its own weight. This compaction reduced the pore space and subsequently caused a portion of the residual interstitial melt to evacuate its shrinking reservoir (Bachmann and Bergantz, 2004). Segregation of fractionated, interstitial melt may also have resulted as a consequence of the destabilization of stagnant mush during recharge. Regardless, the buoyant, high-silica melt would then migrate upward by porous ﬂow or via dikes through the crystal-rich mush. In fact, ﬁeld relations indicate that many pulses of high silica granite were emplaced at the roof as sheets (dikes and sill-like bodies). Zircon samples from this unit yield ages that bracket most of the life span of the Spirit Mountain granite (16.1–17.2 Ma), suggesting that segregation events occurred throughout the assembly of the Spirit Mountain granite. Ascent pathways for these fractionated melts (i.e., feeder dikes and/or intergranular channels) are only rarely preserved within the underlying granite and quartz monzonite, probably because they were disrupted by subsequent destabilization of the host mush or because they collapsed after drainage of melt. The resulting cumulate would be enriched in the earlier crystallizing phases (feldspars, biotite, accessories), and poor in quartz, a late crystallizing phase. This process would plausibly have the most impact on the bottom of the mush pile, where the pressure is the greatest. This can explain the large zone of quartz monzonite at the bottom of the Spirit Mountain granite, which is (1) enriched in feldspars, biotite, and accessories; (2) depleted in quartz; and (3) strongly foliated perpendicular to the up direction. Architecture of the Batholith The composite, multistage constructional architecture of the Spirit Mountain batholith hinted at by the Spirit Mountain granite is made clearer by the younger units. The diorites and the ﬁnegrained granite (FGG) are exposed as hundreds of initially subhorizontal sheets, essentially exhibiting sill-on-sill geometry. We think it plausible that the Spirit Mountain granite formed in a similar manner. Although little evidence for the initial form of these intrusive pulses was preserved, we consider it very likely that they had sheet-like geometry similar to that of well-preserved FGG. These sheets, in our view, subsequently merged as a consequence of mush destabilization, producing a gradational, relatively homogeneous body with, for the most part, only subtle internal contacts.

Spirit Mountain batholith and Secret Pass Canyon volcanic center

17.4 Ma

~17 Ma

~16 Ma

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Figure 5. Cartoon of the assembly of the Spirit Mountain batholith. Drawings by Rob Smith.

Field relations and zircon geochronology discussed above suggest the following history for the Spirit Mountain batholith (Fig. 5): 1. Circa 17.4 Ma: The magma associated with the roof unit was emplaced, followed by a brief pause in magmatism. Zircons associated with this initial phase of emplacement were later redistributed throughout most of the batholith. 2. Circa 17–16 Ma: Sporadic injection and crystallization of granitic magma formed the bulk of the Spirit Mountain granite. Magma was probably emplaced as horizontal sheets within a semi-rigid crystal mush. Injections par-

tially remobilized the surrounding area, entraining zircon antecrysts, and creating (small?) local magma chambers. Within these local chambers, fractional crystallization occurred, producing high-silica melt that migrated upward toward the roof. Minor amounts of maﬁc magma were injected intermittently, becoming small pods and/or maﬁc enclaves. Our U-Pb data cannot conﬁdently establish an order for steps 3–7, as samples have multiple closely spaced apparent ages and our interpretations of emplacement ages fall within error of each other. Field relations and zircon data are consistent with the following sequence and approximate timing:

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Lang et al. 3. Circa 16.0 Ma: The Mirage granite intruded, probably in sheet-by-sheet fashion. Fractional crystallization of this magma produced a leucocratic roof zone, similar to but much smaller than that of the Spirit Mountain granite. 4. Circa 15.9 Ma: Injections of basaltic-dioritic magma formed horizontal sheets and pods that cut the Spirit Mountain granite and were in part coeval with the Mirage granite and FGG. 5. Circa 15.9 Ma: FGG was emplaced as a series of successively stacked horizontal sheets. Cooling of these injections was probably relatively rapid, as suggested by the ﬁne-grained texture. Minor maﬁc to dioritic magma emplacement continued through this time, as well. Field relations are consistent with FGG being at least in part coeval with the Mirage granite (see no. 3). 6. Circa 15.8–15.7 Ma: Magma continued to be injected into the Spirit Mountain granite, which was relatively rigid (low melt fraction) by this time to produce the distinct, younger sequence. Fractional crystallization again produced highsilica melt, which accumulated at the roof zone of this intrusion. Some of this fractionated melt debouched from an intra-batholith cupola upward into the overlying granite. A preserved network of leucogranite dikes, sills, and pods document the ascent pathways and ponding zones of this material. This new magma entrained abundant zircons and presumably other crystals from the extant granite (mush?) and apparently provided enough heat to dissolve zircons and subsequently induce new in situ rim growth. 7. Circa 15.3 Ma: Termination of Spirit Mountain batholith magmatism was marked by injection of a series of vertically intruded, felsic to maﬁc dikes. Emplacement of the Newberry dike swarm followed ﬁnal solidiﬁcation of the remainder of the batholith and may have been facilitated by the onset of rapid E-W extension (George et al., 2005), which is suggested to have commenced ca. 16.0 Ma (Faulds et al., 2001).

The Secret Pass Canyon Volcanic Center Overview On this part of the trip, we explore a portion of the Secret Pass Canyon eruptive center mapped by Lang (2001). The Secret Pass Canyon volcanic center comprises a sequence of intermediate to very felsic volcanic and hypabyssal rocks and volcanogenic sediments that were emplaced between 18.5 and 17.4 Ma. The volcanic rocks are dome-related breccias and ﬂows and airfall and ash ﬂow tuffs. The hypabyssal intrusive rocks include dikes, sills, peperites, a volcanic neck, and, according to our interpretation, a thick series of sills that merges to form a very shallow laccolith. This trip will visit highlights within the 20 km2 area that we have studied. The Secret Pass Canyon volcanic center may be much larger, however. Reconnaissance and previous work indicate that a very similar lithologic package continues and thickens southward for 15 km toward Oatman, Arizona, where it appears

to terminate against a large caldera complex (Thorson, 1971; Liggett, 1981; Lang, 2001). Geologic Units and Field Relations The portion of the Secret Pass Canyon volcanic center that we will examine on this trip resides in a fault block bounded to the west by the east-dipping Arabian Mine Fault and to the east by the west-dipping Frisco Mine Fault (Fig. 6). The local basement, 1.4 Ga Proterozoic granite (unit Xg; Davis Dam Granite of Faulds et al., 2000), is exposed in the footwalls of the block-bounding faults, but not within the studied area of the Secret Pass Canyon volcanic center. The Miocene section within the fault block dips consistently ~30°NE. It records a transition from intermediate to felsic and then back to intermediate magmatism. The sequence, from bottom to top (Fig. 7), consists of a trachydacite (unit Td) that interﬁngers with volcanogenic sediments and tuffs (unit Tst), overlying volcanogenic sediments and tuffs (unit Tvs), rhyolitic breccias (unit Tbr), and lava ﬂows (unit Trf). This sequence was injected by a myriad of intrusives including a rhyolite dome (unit Tir), trachydacite neck, and intermediate to felsic dikes and sills, some of them peperitic (units Ti). Perhaps the most intriguing aspect of the Secret Pass Canyon volcanic center is the thick trachydacite and interﬁngering volcanogenic materials that constitute the lowest exposed part of the section. At ﬁrst glance, these two units, which represent a majority of the map and are >1 km thick (no base exposed), appear as a monotonous gray mass of dissected pediments, rugged hills, and small mountains. Within this monotonous mass, however, is a sequence of resistant and non-resistant rocks that vary in thickness from meters to tens of meters. Field inspection shows that the resistant sections consist of trachydacite, whereas the less-resistant sections consist of volcaniclastic sediments and occasional tuffs. The trachydacite forms ledges that parallel the dip of volcanic and sedimentary strata, but they lack apparent ﬂow tops or bottoms (no vesicle concentrations and very few well-deﬁned rubbly brecciated zones). The trachydacites are porphyritic, with ~20%–40% millimeter to centimeter phenocrysts (mostly plagioclase and biotite, plus altered pseudomorphs; some fresh samples also have cpx). Very sparse maﬁc enclaves ~1–10 cm in diameter can be found within the section. The trachydacites show modest compositional variability, from ~62 to 67 wt% SiO2, and are rich in alkalis (subequal Na2O + K2O total ~8 wt%) and incompatible trace elements. 40 Ar/39Ar (biotite) ages of two trachydacite samples are essentially identical at 18.55 Ma (Table 1). In part because the interﬁngering volcanogenic package Tst is weakly resistant, Tst and Td contacts are poorly exposed, and it is not obvious whether Td represents resistant, thick ﬂows with poorly developed bases and tops interbedded with weaker, dominantly sedimentary strata, or whether Td is a large intrusion. The largest exposure of Tst is ~4 km long and ranges up to >100 m thick. It comprises conglomerates and coarse breccias (probable lahar deposits), rhythmically bedded volcanogenic sands, and a ~5-m-thick welded tuff. These strata have attitudes consistent with those of the ledge-forming trachydacite and the overlying volcaniclastic materials (unit Tvs). Sheets

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Normal fault; ball on downthrown side; solid where known, dashed where approximate

A 30

E

Ex

N

ten fm

to

Ti

Ti

HW

B

g

pin

ap

Tbr

Tvs Td

35°07’30” N

Strike and dip of bedding

Td

Tst

114°25’ E

Tbr

Other Y

A

68 Arizona State Highway 68 Geochron sample locales

Ti 114°22’ E

Figure 6. Simpliﬁed bedrock geologic map of the Secret Pass Canyon volcanic center (SPCVC), located here within a relatively coherent fault block and consists of a >1 km thick sequence that records intermediate to felsic back to intermediate magmatism from ca. 18.5 to 17.4 Ma. Crustal extension between 16 and 11 Ma tilted the Secret Pass Canyon volcanic center ~35°NE, affording a cross-sectional view of the sequence. See Table 1 for more geochron sample details. NW AZ—northwest Arizona.

of trachydacite that interﬁnger with this relatively thick sequence can be interpreted as either sills or ﬂows. The most compelling exposure is at the top of Td in the central part of the map area, where the contact between Td and overlying brown, rhythmically bedded volcanogenic sandstones of Tvs is clear. Approaching this contact from below, discontinuous stringers and angular fragments (centimeter to several decimeter scale) that are lithologically very similar to the Tvs sandstone, but with highly variable orientations, become prominent within Td. Within meters of the contact, angular fragments of Tvs enclosed within Td are common, and centimeter-scale sheets of Td are traceable along bedding planes of the sandstone just above the contact. Td very near the contact is brecciated and altered. We interpret these relations to indicate that the trachydacite intruded rather wet but consolidated sandstone. Fragments of the sandstone in Td indicate microstoping. The fact that

the contact is largely concordant (and highly concordant on map scale) and that it has an orientation consistent with the attitudes throughout both the overlying Tvs and underlying Tst suggests that the roof of the intrusion was horizontal to subhorizontal. This strongly suggests that all of Td is intrusive, that ﬁngers of Td in Tst are sills, and that the Td ledges in otherwise monotonous trachydacite may reﬂect tops of sills that inject sills. Similar stringers and fragments of sandstone exposed in the upper part of Td elsewhere in the study area probably reﬂect the same processes. In sum, this huge sill complex seems to constitute a very ﬂat-topped, shallow sill-on-sill laccolith, perhaps emplaced in a fashion similar to that described in the Henry Mountains (Morgan and Tikoff, 2008). It apparently marked a fresh infusion of magma into the uppermost crust after the volcanic episode indicated by Tst and the period of quiescence marked by the lower Tst sandstones.

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Trf Tbr

Tir

Tbr Tvs

Tst

Tst Tst

Td

Td Tst

Td

Ti

Tst

Xg

Tst

Td Figure 7. Generalized stratigraphic column of the Secret Pass Canyon volcanic center. Tbr—rhyolite breccias; Td—trachydacite; Ti—felsic dikes and sills; Tir—rhyolite dome; Trf—lava ﬂows; Tst— volcaniclastic sediments and tuffs; Tvs—volcaniclastic sediments; Xg—Proterozoic granite.

40

The incompletely exposed sequence mapped as Tvs is ~100– 150 m thick. Above the brown, well-bedded sandstones, it includes airfall tuffs and reworked tuff. The reworked airfall tuffs consist of repeating sequences of red and white, moderately sorted, coarse- to ﬁne-grained sandstones that are 3–10 cm thick. They are primarily made up of 1–2 cm pieces of pumice, possibly reﬂecting Plinianstyle activity leading to emplacement of an overlying ignimbrite and followed by silicic dome emplacement (Cas and Wright, 1987). The 15–20-m-thick ignimbrite is a rhyolitic, sanidine-bearing welded tuff with a vitrophyric base that sits directly above the red and white pumiceous sandstones. It is overlain by a thin (~0.5 m), poorly consolidated tuff or tuffaceous sandstone. The ignimbrite yields a 40Ar/39Ar sanidine age of 17.48 Ma (Table 1). The age of the ignimbrite raises interesting questions about the validity of our laccolith interpretation of Td and, if the interpretation is retained, places important constraints on laccolith emplacement in time and space. If the intrusive relation between the trachydacite and the andesitic sandstones is correct, then the sandstone must be at least 18.55 Ma, meaning that a 1 m.y. period of time is reﬂected in unit Tvs. If the reworked airfall tuffs do reﬂect Plinian-style activity, then it seems logical to think that the red and white reworked tuffs were emplaced close in time to the 17.5 Ma ignimbrite. If true, this indicates that a 1 m.y. hiatus exists between the base of the reworked tuffs and the top of the andesitic sandstone, and it strongly suggest that the top of the laccolith was extremely shallow—perhaps on the order of 100 m or less. This is entirely possible, because the section is not continuously exposed, but we have found no direct evidence for the implied disconformity. Dating of the volcanic fragments in the sandstone fragments may help resolve this uncertainty. Deposited on top of the ignimbrite and overlying tuff are massive, interﬁngering rhyolite breccias and lava ﬂows. The breccias (unit Tbr) form striking cliffs hundreds of meters tall that consist of multiple individual units. Although lumped together as one unit here, unit Tbr is divisible into distinctive avalanche breccias and block and ash ﬂows, each interpreted to be associated with a different style of silicic dome collapse. Avalanche breccias are predominantly clast-supported and are composed of pieces of reworked ﬂow banded rhyolite that range from ~5 cm to 1 m in size. Near the base of unit Tbr, avalanche breccia deposits

39

TABLE 1. SUMMARY OF Ar/ Ar DATES FOR THE SECRET PASS CANYON VOLCANIC CENTER Sample no.

Rock type

Map unit

Lat/Long

5/9-#5

Trachydacite

Td

5/11-#2

Trachydacite

Ti

5/14-#3

High-silica rhyolite

Tbr

5/29-#2b

High-silica rhyolite

Tbr

Trachydacite

Td

Ignimbrite

Tvs

N35° 8.61′ W114° 24.08′ N35° 8.06′ W114° 22.97′ N35° 9.16′ W114° 23.10′ N35° 9.99′ W114° 23.95′ N35° 8.47′ W114° 24.59′ N35° 8.95′ W114° 24.33′

1/4-#4 5/11/01-#3

Material dated Biotite

Age (Ma) 18.56 ± 0.06

Map indication A

Biotite

17.45 ± 0.07

B

Sanidine

17.34 ± 0.06

C

Biotite

17.68 ± 0.08

D

Biotite

18.55 ± 0.09

E

Sanidine

17.48 ± 0.1

F

Spirit Mountain batholith and Secret Pass Canyon volcanic center

Time 1

Time 3

197

Time 2

Time 4

Figure 8. Cartoons illustrating the geologic history of the Secret Pass Canyon volcanic center. Time 1 (>18.56 Ma): Intermediate volcanism that erupted ignimbrites and possible lavas with coeval lahar and volcanogenic sediment deposition (dark gray layers and unit Tst in Fig. 6), all on Proterozoic granite basement. This was followed by a period of quiescence as rhythmically bedded sandstones (white stippled layers) that form the base of unit Tvs accumulated in a depositional basin. Time 2 (18.56 Ma): Rejuvenation of intermediate magmatism that resulted in the emplacement of trachydacite sills that intruded (and inﬂated) the earlier emplaced layered deposits. It is unclear if sill emplacement was associated with any eruptions. Time 3 (ca. 17.5–17.4 Ma): Eruption of ignimbrites and rhyolitic domes as well as rhyolite lava ﬂows. Time 4 (ca. 17.4 Ma): Abrupt transition to intermediate magmatism marked by volcanic necks, dikes, and sills.

also include clasts of trachyandesite. In contrast, the block and ash ﬂows consist of rhyolite clasts up to 20 cm in size, supported in a matrix dominated by ash and unﬂattened pumice. 40Ar/39Ar dating of rhyolite clasts from a block-and-ash-ﬂow deposit and a rhyolite ﬂow yield ages of 17.68 Ma (biotite) and 17.33 Ma (sanidine), respectively. SiO2 concentration in clasts from the breccia deposits range from 68 wt% to 73 wt%. Samples from the rhyolite lava ﬂows (Trf) overlap in composition with the breccias but tend to be more silicic, ranging from 72 wt% to 82 wt% SiO2 (we infer that the most silicic samples have been altered, but some are probably true high-silica rhyolites). Numerous intrusions cross-cut units Tvs, Tbr, and Trf. These include a rhyolite dome (unit Tir), volcanic necks of trachydacite (unit Tid), and intermediate and felsic dikes and sills (unit Ti). The rhyolite dome is in the northwest corner of the map area, and the trachydacite necks (unit Ti) are to the southeast. The 40Ar/39Ar (biotite) age of one the necks is 17.45 Ma; thus, although it intrudes Tbr, it was essentially contemporaneous with the rhyolitic volcanism. For simplicity, the maﬁc to felsic dikes are grouped together. Rhyolite dikes are exposed only where they intrude Proterozoic granite (unit Xg), outside the fault block that is our focus. Dikes of

intermediate composition are aligned NW-SE, trending toward the trachydacite volcanic necks (Fig. 6), suggesting that they may all be part of a common feeder system. Where these dikes and related sills intrude into unit Tbr, they break up in peperitic fashion. This suggests that the breccias were wet, and probably young, when they were intruded by somewhat more maﬁc magmas. History of the Secret Pass Canyon Volcanic Center Based on the examination of ﬁeld relations and petrographic, geochemical, and geochronologic data (Table 1) summarized above, we suggest the following history for the Secret Pass Canyon volcanic center (Fig. 8): 1. ≥18.55 Ma: Intermediate to felsic volcanism that erupted ignimbrites and lavas with coeval deposition of lahars and volcanogenic sediments (Tst). This was followed by a period of quiescence, as suggested by the rhythmically bedded brown sandstones at the base of unit Tvs. 2. 18.55 Ma: Emplacement of numerous trachydacite sills into Tst and the base of Tvs. The sills are interleaved and merge into a massive laccolith. No direct evidence that these trachydacite magmas reached the surface has been identiﬁed.

114º 44’ 53.14” W

35º 18’ 53.43” N

198

114º 36’ 2.32” W

Lang et al.

A

35º 08’ 14.87” N

35°12’08” N

B 1 kilometer

Arizona HWY 68

1 mile

N

1 35°10’ N

2

3a

3

35°07’30” N

3b

114°25’ W

114°22’ W

Figure 9. Maps showing ﬁeld trip stop locations. (A) Spirit Mountain Batholith (Day 1; taken from Google Earth). (B) Secret Pass Canyon volcanic center (Day 2). Dashed line with arrows in (B) refers to the walking route we will take.

Spirit Mountain batholith and Secret Pass Canyon volcanic center 3. From 18.55 to ca. 17.7 Ma: A possible period of quiescence within the volcanic center. This time may mark a period of magmatic differentiation from intermediate to felsic compositions. 4. Circa 17.7–17.4 Ma: Felsic eruptions: ignimbrites, airfall, dome formation, and collapse. 5. Circa 17.4 Ma: An abrupt transition to intermediate magmatism as marked by the volcanic necks as well as dikes and sills. FIELD TRIP STOPS (Fig. 9) Day 1: Spirit Mountain Batholith Mi Cumulative Description 0.0

–

3.4 14.5

3.4 17.9

50.3

68.2

7.5

75.7

0.1

75.8

Leave hotel from the University of Nevada–Las Vegas; get on Tropicana, and start mileage. Merge onto 515/U.S. 93 (south). Follow the interstate out of Las Vegas and then merge onto U.S. 95 south. Turn left onto Christmas Tree Pass Road (east). Turn right onto dirt road heading up the hill (south). Stop 1: NO HAMMERS AT THIS STOP— this is sacred Native American land.

Stop 1: Spirit Mountain Batholith Roof We are now in the roof of the Spirit Mountain batholith, which is a ~25 km × 2 km zone of high-silica leucogranite. This zone comprises sheets of aplite, porphyry, and ﬁne- to mediumgrained equigranular granite, with contacts that are sharp to barely perceptible. Most sheets were initially subhorizontal, but dikes (subvertical) are also common. We interpret these relations to indicate repeated emplacement of the leucogranites—some sheets intruding a hot, melt-bearing mush, and some intruding solid rock. Vesicles, or miarolitic cavities, are widespread and most common toward the west (top). Pegmatite pods and dikes dominated by coarse quartz and alkali feldspar are present locally. Typical leucogranites have ~40%–50% alkali feldspar (subhedral), 30%–40% quartz (anhedral in groundmass, but phenocrysts are subhedral), ~10% plagioclase (sub-euhedral), and ~1% biotite (euhedral). Porphyritic variants contain ~0.5 cm phenocrysts of quartz and alkali feldspar.

0.4

80.0

1.6

81.6

Here you will see a large, rounded outcrop to your right (south); this is loosely known as the Catacombs. The road turns right (south), and you follow. Stop 2: NO HAMMERS AT THIS STOP; STILL SACRED LAND.

Stop 2: The Sacatone Wash–Quartz Monzonite Cumulate From the parking area, walk up the wash (west) for ~0.2 mi. The “wash” at this point turns into a steep-walled canyon. Continue on for ~0.1 mi to see good exposure of the quartz monzonite and giant maﬁc enclaves. The leucogranite from Stop 1 grades downward (eastward) through granite into magmatically foliated quartz monzonite, seen here, that is poorer in quartz and richer in biotite. The quartz monzonite is coarse-grained, with 40%–50% alkali feldspar (euhedral), 30%–35% plagioclase (euhedral), 10%–15% biotite (euhedral), 5%–15% quartz (interstitial, anhedral), and ~1%–2% hornblende (subhedral). Foliation, deﬁned by aligned biotite, alkali feldspar, and plagioclase, is parallel to paleohorizontal and gradually becomes stronger downward. Dioritic enclaves are abundant, very large (up to 3m), pancake shaped, and oriented parallel to the rock’s ~N-S striking, W-dipping fabric (Fig. 10). Based on euhedral to subhedral feldspar crystal shapes and weakly to unstrained interstitial quartz, this fabric is interpreted to be dominantly magmatic and probably related to compaction of a crystal mush (cf. Bachmann and Bergantz, 2004). Mi Cumulative Description 3.0

84.6

4.5

89.1

Get back on Christmas Tree Pass Road and drive south until the road dead ends into NV-163. Turn right (west) onto the highway. Follow the highway. Turn right (north) on a dirt road, which is across the highway

Mi Cumulative Description 0.1

75.9

3.7

79.6

Turn around and return to Christmas Tree Pass Road. Turn right (east). The dark colored outcrops to the left (north) are Precambrian country rock, consisting of megacrystic granite and orthogneiss.

199

Figure 10. Dioritic enclaves within the leucogranite.

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Lang et al.

0.1 0.5

89.2 89.7

1.0

90.7

from a large, sandy area used for off-road vehicle recreation. Road splits; stay right. Road crosses underneath a power line and across a power-line road. Continue straight. Continue straight on the road, keeping right if possible. This road can be fairly bad, so a 4WD truck is recommended. The road then turns sharply to the left (west). Don’t continue on, but park here. This is Bridge Canyon.

Stop 3: Bridge Canyon We are overlooking Bridge Canyon. Walk ~0.3 mi north, downhill into the canyon. Once in the canyon, go right (east), “downstream.” Here we see the granitic (sensu stricto) portion of the Spirit Mountain granite, some felsic dikes, and a few large pods of high silica granite segregate. The granite here is fairly homogeneous and is characterized by 30%–40% pink alkali feldspar (sub-euhedral), ~20%–35% plagioclase (euhedral laths), 15%–30% quartz (interstitial and anhedral to ~1 cm subhedral), 3%–8% biotite (euhedral), and ~1% sphene (euhedral). At this location, quartz forms prominent, round grains, but its abundance decreases markedly to the east. Locally, small (<2 cm), plagioclase + ﬁne-grained biotite-rich clusters are present within the granite. Fine-grained dioritic enclaves are present here and throughout the lower ~1/2 of the coarse granite. They are typically ellipsoidal, from ~5–30 cm, with irregular margins that are penetrated by crystals of the host granite, suggesting liquid-crystal mush (or mush-mush) contact. The enclaves are composed mainly of plagioclase and biotite, with minor hornblende and clinopyroxene. Some contain large alkali feldspars, suggesting crystal incorporation from the host granite. These enclaves are the only manifestation of maﬁc input during solidiﬁcation of the Spirit Mountain granite unit. Widespread schlieren show no apparent preferred orientation. Abundant pegmatite pods are commonly bounded by schlieren at their paleo-upper surfaces. Walk ~0.5 mi down canyon. At about this point, a distinct intrusion marks a break in the “homogeneous gradation” within the Spirit Mountain granite. In the canyon, the contact is very subtle due to the similarity of the rock, so it is best to walk south out of the canyon. This intrusion also grades from a high silica leucogranite cap (1–100 m thick) downward into coarser, less felsic granite. The contact between this unit and the overlying granite ranges from straight to very sinuous. In places, large vesicles and pegmatites are common in the leucogranite below the contact. Locally, small (<1 m) blocks of the overlying granite are also present just below the contact. About 0.5 mi south of the canyon, we encounter the cap of this distinct intrusion that forms a rampart dubbed “Jabba the Hut” (Fig. 11), which is more resistant to weathering probably because of its quartz-rich mineralogy. A network of leucogranite dikes and sills emanates from the top of this intrusive sequence into

the overlying granite. Here we also see large pods of high-silica leucogranite, interpreted to be differentiated magma that originated from the underlying intrusion. Pods are fairly large—up to 500 m long × ~150 m thick—and elongated in the paleohorizontal direction and are bounded by sharp contacts on all sides. Felsic dikes are also present at this stop. The dikes were originally oriented vertically and are generally porphyritic, with round quartz phenocrysts and varying amounts of feldspar phenocrysts as well. The groundmass is gray and quenched in places against the Spirit Mountain granite. The dikes seem to be the favored canvas on which petroglyphs were carved, probably due to their typical smooth weathering surfaces and dark brown desert varnish. No hammering near these artifacts, please. Mi Cumulative Description 1.6

92.3

1.3

93.6

0.5

94.1

Turn around and get back to Hwy 163. Turn left (east). Drive on the interstate until just before the hard turn left (north). Just before that turn, turn off the road to the right. Watch the guard rails—just before they begin, that’s where you turn right, off the road. (A 4WD vehicle is recommended for the next leg.) From the highway, a network of dirt roads heads east. Get on one of these roads and drive as far as you can (~0.5 mi). About 0.4 mi in, there is a steep stream bed. Cross this at your own risk. (Sometimes it’s passable, sometimes it’s not.) If you can, drive for another 0.1 mi, and park near a building foundation. This is Hiko Canyon, and there should be a large, oddly developed Cottonwood tree within sight.

Stop 4: Hiko Canyon At this stop, we will traverse part of the lower exposed portion of the Spirit Mountain batholith, which documents distinct, later injections into the coarse-grained Spirit Mountain quartz

Figure 11. “Jabba the Hut”: an outcrop of high-silica leucogranite.

Spirit Mountain batholith and Secret Pass Canyon volcanic center

201

monzonite. We will be walking east and thus down-section in the Spirit Mountain batholith. New units we encounter include the ﬁne-grained granite and the diorite, both of which intrude the Spirit Mountain quartz monzonite cumulate, which is also seen in abundance at this stop. The goal is to walk ~0.7 mi down the canyon to where there is an excellent exposure of the interaction of these three units, with mechanical and apparently some diffusive mixing and disaggregating blocks of the older cumulates in the newer magma pulses. Along the way, we will see evidence of the ﬁne-grained granite intruding the Spirit Mountain quartz monzonite. Begin by walking ~50 m to the east from the parking area; you will encounter a large sheet of rock, weathered brown, with numerous petroglyphs. It is a planar body, dipping to the west. This is the ﬁne-grained granite. Stop 4a: Fine-Grained Granite The ﬁne-grained granite (FGG) is uniformly ﬁne-grained and equigranular, except for the local alkali feldspar megacrysts that are present in some sheets. It typically has 30%–35% plagioclase (small, euhedral laths), 25%–30% quartz (anhedral), 30%–35% alkali feldspar (blebby, anhedral, or megacrysts), and 5%–10% biotite (euhedral). Myrmekitic intergrowths of quartz and alkali feldspar are common. Megacrysts of alkali feldspar, where present, have slightly irregular and inclusion-rich boundaries, perhaps indicating minor resorption followed by regrowth. At many locations, biotite is aligned and the quartz is strongly strained with conspicuous subgrain development, suggesting minor to strong subsolidus deformation of the FGG unit. The FGG is observed primarily as west-dipping (initially subhorizontal) sheets to pod-like intrusions within the middle to deeper parts of the southern Spirit Mountain granite. Sharp cross-cutting relationships clearly show that FGG intruded the Spirit Mountain granite. Blocks of the Spirit Mountain granite (in some cases disaggregated or in situ screens), ranging in size from centimeters to >100 m, are commonly seen within or in contact with the FGG. Individual FGG sheets, where discernible, are centimeters to 50 m thick, though in many places it is difﬁcult to identify separate sheets. In some areas, internal contacts separate FGG phases that are distinguishable by the presence or absence of 1–2 cm alkali feldspar megacrysts. It is unclear whether these megacrysts were derived from the Spirit Mountain granite host rock or if they grew from the FGG magma. About ~0.25 mi in, the canyon turns south brieﬂy, and exposed high on the west-facing wall is a large Spirit Mountain granite block completely engulfed by the younger ﬁne-grained granite. Stop 4b: Diorite Relatively maﬁc rocks, typically dioritic, are exposed as initially subhorizontal sheets to pod-like intrusions up to ~100 m thick. The diorite is ﬁne- to medium-grained, and typically contains close to 50% hornblende, ≥50% plagioclase, <5% quartz, and up to 5% biotite. Sphene, apatite, opaque minerals, and

Figure 12. Blobby pillows of diorite in a contaminated ﬁne-grained granite.

minor zircon are present as well. Locally, these sheets pinch out and then reappear along strike. In places, the diorite is also present as pillows in the ﬁne-grained granite unit (see above). Though visually striking, the diorite is, volumetrically, a very minor unit of the Spirit Mountain batholith. About 0.5 mi down the canyon, blobby pillows of diorite within the FGG (Fig. 12) indicate that intrusion of the two magmas coincided temporally, but angular enclaves of diorite in FGG and sharp contacts of FGG dikes into diorite indicate that at least some of the FGG intruded after diorite became rigid. This, along with the internal contacts, suggests that there were multiple pulses of FGG emplaced into the Spirit Mountain batholith. Small dikes of the FGG cross cut portions of the Mirage granite, although a clear contact between the two units has not been observed. About 0.6 mi down the canyon, a large exposure on the north (left) wall of the canyon illustrates the sheet-on-sheet nature of these later intrusions (Fig. 13). Closer inspection suggests that these FGG sheets intruded the larger diorite mass. Continue ~50 m down the canyon, and on the south side of the stream you will encounter a large, polished outcrop displaying many complex relationships between the FGG, the diorite, and the Spirit Mountain quartz monzonite. The outcrop is mostly a ﬁne-grained rock. It appears to be dominantly diorite, but rocks in varying shades of gray cut or grade into one another, possibly suggesting hybridization between FGG and diorite. The ﬁne-grained host is littered with xenoliths, xenocrysts, and small dikes. The xenoliths include Spirit Mountain quartz monzonite, and, rarely, dark megacrystic Precambrian granite. Whereas some of these blocks are angular and have sharp margins, others have irregular, poorly deﬁned margins. In places, the host maﬁc rock is contaminated by feldspar crystals, suggesting partial disaggregation of the granite blocks. At this stop, many xenoliths can be seen in what appears to be a spectrum of disaggregation, from angular blocks to loose collections of feldspars within contaminated diorite (Fig. 14).

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up

felsic

mafic

felsic

mafic

felsic

mafic

Figure 13. Successively stacked ﬁnegrained granite (FGG) sills intruded into diorite. This sill-on-sill geometry is observed throughout the FGG and diorite units. For scale, the FGG sills are ~5 m thick. Initial up-direction is indicated.

Mi Cumulative Description 0.5

94.6

Return to the highway via the dirt roads.

Day 2: Secret Pass Canyon Volcanic Center All stops on the trip are on the Union Pass, Arizona, 7.5 min quadrangle. Mi Cumulative Description 0.0

–

0.3

0.3

7.2

7.5

0.8

8.3

0.1

8.4

0.4

8.8

Drive north on S. Casino Drive in Laughlin, Nevada, turn right (east) onto Highway 163, and cross the Colorado River. Turn left (north) onto Hwy 95 (Mohave Valley Highway); it will become Hwy 68 when the road makes a big bend to the east. Turn right (south) onto an unnamed road. Go up the hill and to the left (southeast); be careful on this road as it is traveled by large trucks hauling gravel from a local quarry (along this road, we are driving through the local basement of Proterozoic granite [unit Xg]; the dark brown resistant ridges on either side are rhyolite dikes that intruded the basement and may have fed the Secret Pass Canyon volcanic center volcanic system). The road splits within a wash—go to the right (southwest). The road splits again—keep to the left (south). On the right (southwest) is a road that goes up a small hill. Go up to the top of the hill and park.

Stop 1: Overlook of the Secret Pass Canyon Volcanic Center The road stops at the top of the hill (35°10′30.75′′ N, 114°26′41.72′′ W) providing a panoramic view of the Secret Pass Canyon volcanic center. We are standing on altered Td

Figure 14. A xenolith of Spirit Mountain granite within the diorite. Poorly deﬁned margins of the xenolith and large feldspar crystals in the diorite suggest that this xenolith was partially disaggregated.

(trachydacite). Directly to the north is an intrusive rhyolite dome (unit Tir). The dome is highly altered, as indicated by its deep orange color and very high SiO2 (>80% SiO2). At this locality, the dome has a planar contact with unit Td; however, to the north, the dome has locally overturned units Tvs (volcaniclastic sediments) and Td (Fig. 15A). To the east of the dome is a thin spire called Thumb Butte (Fig. 15B), which is composed of a mixture of avalanche and block-and-ash-ﬂow deposits of unit Tbr that we interpret as emplaced during a cycle of rhyolitic dome formation and collapse. The breccias at the base of Thumb Butte contain abundant clasts of ﬁne-grained trachydacite, but felsic clasts become dominant farther up the butte and comprise most of the breccia units within the Secret Pass Canyon volcanic center. To the east of Thumb Butte is a crescent-shaped basin and ridge (Fig. 15B). The bottom of the basin consists of the volcanogenic materials comprising unit Tvs. Above unit Tvs and comprising the steep cliffs within the ridge surrounding the basin are the intermixed breccias observed at Thumb Butte. Capping the ridge are rhyolite ﬂows of unit Trf. The rhyolite ﬂows are resistant to erosion and cap the tops

Spirit Mountain batholith and Secret Pass Canyon volcanic center

203

A

Intrusive rhyolite dome (unit Tir)

Overturned trachydacite (unit Td) and volcaniclastic materials (unit Tvs)

B Thumb Butte Crescent-shaped basin and ridge

C Trachydacite volcanic neck (unit Ti)

Trachydacite (unit Td) Figure 15. Sites from the panoramic view from Stop 1. (A) Intrusive rhyolite dome (unit Tir). (B) Thumb Butte and the crescent shaped basin and ridge. (C) The volcanic neck of trachydacite (unit Tid). (Continued on following page.)

204

D

Lang et al.

E Spirit Mountain and Spirit Mountain Batholith Newberry Detachment Fault

Lower Plate Upper Plate

Figure 15 (continued). (D) View of the monotonous nature of the trachydacite sill complex including the possible trachydacite dome(s) (unit Td). (E) Distant view of the Newberry Mountains including Spirit Mountain and the Spirit Mountain batholith. The east-dipping Newberry detachment fault marks the base of the Newberry Mountains.

Rhyolite flow (Trf )

Volcaniclastics (Tvs) Ti (intermediate sills)

Breccias (Tbr)

Breccias (Tbr)

Volcaniclastics (Tvs) Trachydacite sill complex (Td)

Figure 16. Stratigraphy of the Secret Pass Canyon volcanic center as seen in a mountainside at Stop 2. The trachydacite complex (unit Td) comprises the low hills in front of the mountain. The bedded nonresistant materials are volcaniclastic sediments of unit Tvs. The breccias of unit Tbr make the two cliff-forming units in the middle of the mountain, and a rhyolite ﬂow of unit Trf makes the topmost point on the mountain. The dark bands outlined in a white dashed line are sills that have intruded into the breccias.

Spirit Mountain batholith and Secret Pass Canyon volcanic center of hills where present; the ﬂows at the top of the ridge are also fairly altered (>80% SiO2). Far to the southeast, you will see a sharp pyramidal peak that is surrounded by a blanket of scree (Fig. 15C). This peak is a volcanic neck of trachydacite (unit Ti) and is the collection location of geochron sample B on the geologic map (Fig. 6). Directly to the south are the rugged hills and dissected pediments of trachydacite that we interpret as a thick sill complex or laccolith (unit Td; Fig. 15D). Notice how monotonous this unit appears; we have not yet visited the highest points immediately south of the volcanic neck (unit Ti), but they appear to be continuous with the trachydacite complex. To the west across the Colorado River are Spirit Mountain and the Spirit Mountain Batholith, where we spent yesterday (Fig. 15E). At the base of the mountains is the east-dipping Newberry Detachment fault. You are currently standing on the upper plate of the detachment. Post-volcanic extension has tilted the Secret Pass Canyon volcanic center ~30°NE. At the bottom of the hill, turn left (north), retracing your route.

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Tbr Tbr Ti

Ti Ti Ti

Tbr Ti

Mi Cumulative Description 0.3

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1.9

11

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Turn left (ﬁrst left). We are driving along the east-dipping Arabian Mine Fault (measurements of the Proterozoic granite-Trachydacite contact indicate that the fault dips ~60° E); on the right (west) is Proterozoic granite (unit Xg), and on the left (east) is trachydacite (unit Td)]. Stay left on this road. Turn left (east) and follow the road up the wash (35°09′24.92′′ N, 114°27′40.77′′ W). During this drive, we pass through the trachydacite (unit Td), which again appears as a monotonous mass. The trachydacite continues to the south (right) until it is covered by Quaternary alluvium; reconnaissance to the south has not been able to constrain the bottom of this unit. Intersect with a north-south–trending road (35°9′22.08′′ N, 114°25′21.44′′ W). Our route to the south along this road, toward Oatman, Arizona, is hard to see—the road curves around the east side of a large outcrop of unit Td.

Stop 2: View of the Secret Pass Canyon Volcanic Center Stratigraphy To the north of this junction is a stair-stepped, ﬂat-topped mountain that affords another view of the Secret Pass Canyon volcanic center stratigraphy (Fig. 16). The stratigraphy in this mountain is the same as seen within the crescent-shaped basin and ridge at Stop 1; the sun angle from about mid-afternoon to sunset highlights this stratigraphy magniﬁcently. At the base of the mountain, the small rugged mounds mark the trachydacite (unit Td). The layered nonresistant section is the volcaniclastic sediments and reworked tuffs of unit Tvs. The breccias of unit

Figure 17. Photo of a maﬁc sill (unit Ti) that broke apart as it intruded into presumably wet breccias (unit Tbr). The sill broke apart into pillow-like features outlined with a white dashed line.

Tbr comprise the cliff-forming units within the middle, and a rhyolite ﬂow of unit Trf marks the topmost point. Dark horizontal bands are exposed within the side of the mountain (in between units Tvs and Tbr and within unit Tbr) and represent more maﬁc sills. The topmost sill is broken apart within the breccias, suggesting intrusion into wet volcanogenic material, resulting in explosion features similar to a peperites (Fig. 17). If the sills did intrude into wet, unsolidiﬁed material, then it suggests that there was intermediate to maﬁc input into the magmatic system during or immediately subsequent to felsic volcanism. At this road junction, we are standing within the trachydacite (unit Td). Morphologically, the trachydacite has eroded into dissected hills that seem to have an almost a uniform slope of ~30°NE. Vertical jointing is also strongly apparent and, together with the uniform NE slopes, gives the impression that the trachydacite at this stop represents a suite of lava ﬂows. However, we have found a distinct absence of obvious structures deﬁning ﬂows, such as basal brecciation, rubbly ﬂow tops, and vesicles. Mi Cumulative Description 0.0

13.5

1.1

14.6

Turn right (south; the road from this point on is extremely rough—4WD is necessary!). Along this road, we are completely within the trachydacite complex (unit Td). Pull off within the large wash and park (35°08′42.11′′ N, 114°25′25.38′′ W). As a point of reference, the peak on the east side of the road is labeled on the topographic map as having an elevation of 2540 ft.

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Stop 3: Road to Oatman, Arizona, at Peak 2540 We will be hiking within a wilderness area, so please proceed appropriately. At this stop, we will traverse up-section (east-northeast) through the trachydacite (unit Td) to what we interpret as an intrusive contact with unit Tvs; along the way, we will examine evidence that led to our interpretation of unit Td as an intrusive complex. We will then hike southeast within unit Tvs to the entrance to Secret Pass Canyon, where we will examine the spectacular cliffs comprised of avalanche breccias and block and ash ﬂow deposits. We will then proceed southwest back down-section via Secret Pass Canyon wash. The following is a rough “walking log”; locations that we will spend signiﬁcant time examining are labeled as stops 3a and 3b.

Figure 18. Fresh-cut hand sample of the trachydacite (unit Td) showing its coarse-grained nature. Phenocrysts comprise ~25%–35% of the sample and include plagioclase and biotite. Scale bar is 1 cm.

Trachydacite (Td)

Interfingered sediments (Tst)

Figure 19. Photo showing the general characteristics of the interﬁngered sediments that comprise unit Tst. The materials shown in this photo consist of a series of gray beds several centimeters thick composed of moderately to poorly sorted sand-sized grains of plagioclase, fresh-looking biotite, and lithics.

Mi Cumulative Description 0.0

–

0.0

~0.3

–

~0.3

~0.1

~0.4

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0.8

(35°08′42.11′′ N, 114°25′25.38′′ W): We are beginning within the trachydacite section (unit Td) (Fig. 18). Plagioclase has weathered out in parts of the outcrop surface, giving the false appearance that the trachydacite contains vesicles; in fact, our inspection to date has yielded no vesicles throughout this section. As we walk up the wash, note on the lefthand side, close to the bottom of the outcrop, local lenses of rounded to subrounded trachydacite clasts that are on the order of centimeters in diameter. Compositionally, these clasts appear similar to their trachydacite host, which again appears to be massive and coarse-grained but are distinguishable from the host by slightly different weathering. These clasts represent either (1) material that was picked up during emplacement of the trachydacite, or (2) a brecciated ﬂow base for the trachydacite. (35°08′50.37′′ N, 114°25′13.60′′ W): The wash splits here. Go to the right. (35°08′52.40′′ N, 114°25′5.95′′ W): The wash splits again. Go to the left. In this stretch, interﬁngered with trachydacite, we see volcanogenic sedimentary deposits. These sedimentary deposits deﬁne unit Tst in Figure 6 and outcrop in the walls of the wash as ~2–3-m-thick exposures that are capped by the trachydacite, which has slopes of ~30° NE. These units are gray in color, bedded on 1–10 cm scale. They contain moderately to poorly sorted grains of plagioclase, fresh-looking biotite, and lithic clasts up to tens of centimeters (Fig. 19); the fresh-looking biotite is easily distinguishable in hand samples and suggests that these deposits are locally derived. Several hundred meters farther up the wash (N35°08′52.26′′, 114°24′54.78′′W), we observe a poorly sorted breccia composed of subrounded to angular trachydacite and felsic (rhyolitic?) clasts up to >1 m in size (Fig. 20) that we interpret as a lahar deposit. Continuing up the wash, the volcanogenic deposits are again wellbedded, with fresh-looking biotite. It is worth noting that sedimentary deposits are not the only components of unit Tst. A welded tuff (possible rheomorphic ignimbrite?) is also exposed at 35°08′48.36′′ N, 114° 25′3.96′′ W.

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A

B

C

Figure 20. (A–C) Photos of the possible lahar deposits within unit Tst. The deposits shown in A–C are poorly sorted and composed of subrounded to angular trachydacite and felsic (rhyolitic?) clasts up to >1 m in size.

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Resistant ledge dipping ~30° NE

One trachydacite sill

Figure 21. Photo showing the typical slope characteristics of unit Td. As a whole, ridges that are capped by trachydacite appear to form laterally continuous layers that consistently dip ~30°NE.

A second trachydacite sill

~0.8

~1.3

We are now back into the trachydacite. The trachydacite here is texturally, mineralogically, and compositionally similar to the trachydacite observed at the beginning of this traverse. If you climb on top of one of the walls of the wash, you begin to gain a sense that the trachydacites that cap all of the hills within this section have consistent ENE dip slopes. Further, it appears that, if not for erosion, the caps form laterally continuous layers traceable across much of unit Td (Fig. 21). These layers vary in their thickness from about meters to tens of meters. ~1.3 1.4 This is more of the trachydacite; however, blocks of bedded brown sedimentary material encased within the trachydacite become common in this stretch. These blocks are centimeters thick and vary from centimeters up to about a meter long with varying orientations within the trachydacite; the layers appear to have broken off along distinct bedding planes. Most appear as coherent blocks, but some appear to be warped, suggesting they have experienced ductile deformation. Stop 3a: Units Td-Tvs contact – 1.4 mi (35°9′5.76′′ N, 114°24′23.28′′ W): This is the contact between units Td and Tvs. At the base of the walls of the wash is trachydacite, which is ultimately overlain by rhythmically bedded brown sandstones that comprise the base of unit Tvs. The contact here, although roughly parallel to visible bedding, is locally irregular, with

ﬁngers of the trachydacite intruding up into the sediments and along sediment bedding planes (Fig. 22). When observed along a bedding plane, the trachydacite appears as a brecciated sliver giving the impression that the trachydacite intruded wet sediment, causing it to break apart in a process akin to peperite formation (Fig. 22a). Where the trachydacite has intruded far enough along a bedding plane, blocks of the sediment have fallen—or been stoped —into the trachydacite. Again, these blocks have orientations that vary from parallel to the attitude of the coherent overlying brown sediments to blocks that are nearly perpendicular to this attitude. In many cases, the stoped blocks retain their primary sedimentary features, including bedding, cross bedding, and cut and ﬁll structures. However, as noted in the previous section of this wash, some blocks have experienced ductile deformation. That blocks of overlying sedimentary material have been stoped into an underlying unit strongly indicates that the contact between units Td and Tvs at this locality is intrusive. The observation that the two units are roughly parallel to each other suggests that the intrusion was initially horizontal to subhorizontal— or sill-like—in nature. Based on the observations at this outcrop, together with the massive and coarse-grained nature of the trachydacite lower in the section and the absence of features deﬁning ﬂows, we infer an intrusive origin for unit Td. In this intrusive scenario, the volcanogenic sedimentary deposits observed lower in the section (Tst) are the remnants of the host rock into which the Td intruded. If this interpretation is correct, then the fact that unit Tst has an orientation consistent with the average attitudes at

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A

Brown SS

Td Brown SS Td

Brown SS

Td

Brown SS

B Brown SS

Td Td

Brown SS (incorporated blocks)

Td

Figure 22. (A–B) Photos of what we interpret as an intrusive contact between units Td and Tvs. (A) Fingers of trachydacite that have intruded along individual bedding planes of the overlying brown sandstone have brecciated, giving them the appearance of peperites. (B) Blocks of the brown sandstone that have been stoped into the underlying trachydacite. Individual sandstone blocks have been rotated such that their orientations contrast with those of the overlying sediments. SS—sandstone; Td—trachydacite.

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the Secret Pass Canyon volcanic center suggest that these other intrusions were also sill-like. A test of this hypothesis would be dating the interﬁngered sedimentary deposits. If they yield ages >ca. 18.55 Ma (trachydacites are 18.55 Ma; Table 1), then the intrusive argument is strengthened. Mi Cumulative Description 1.4

~2.5

Climb out of the wash and hike to the southeast (paralleling the cliff face) toward the Secret Pass Canyon wash. The resistant cliffs to the east (left) are the breccias of unit Tbr. The less resistant layers below unit Tbr is the upper part of unit Tvs, which is composed of reworked airfall tuffs that are capped by an ignimbrite (Fig. 23). The reworked airfall tuffs consist of repeating sequences of red and white, moderately sorted, coarse- to ﬁne-grained sandstones that are 3–10 cm thick and contain 1–2 cm diameter pieces of pumice (Fig. 23A), possibly reﬂecting Plinian-style activity here, which may have been a precursor to the eruption of the ignimbrite (Cas and Wright, 1987). The local ignimbrite unit is a 15–20-m-thick, pinkish, sanidine-bearing welded tuff with minor (<1%) biotite (Fig. 23B). It contains a light tan basal vitrophyre that rests directly on the red and white sandstones. It is overlain by a thin (~0.5 m thick), poorly consolidated tuff or tuffaceous sandstone. 40Ar/ 39Ar dating of sanidine within the tuff gives an age of 17.48 Ma (Table 1), raising the question of how much time is represented by unit Tvs. If the trachydacite does in fact intrude the brown sandstone, then this sandstone must be at least 18.55 Ma, meaning that unit Tvs was deposited over an interval of at least 1 m.y. If the reworked airfall tuffs do reﬂect Plinian-style activity, then it seems reasonable that the red and white reworked tuffs were emplaced close in time to the 17.5 Ma ignimbrite. If true, this indicates that a 1 m.y. hiatus exists between the base of the reworked tuffs and the top of the brown sandstone (base of unit Tvs).

Stop 3b: Secret Pass Canyon ~2.5 ~2.7 This big wash is the Secret Pass Canyon wash. Follow this wash northeast into the canyon. As the canyon begins to narrow and the walls become steeper, you get a closer look at the breccias that comprise unit Tbr; individual breccia units are nearly impos-

~2.7

4.1

~4.1

4.9

sible to map within this unit, though some appear to be stratiﬁed (Fig. 24). Rhyolitic lava ﬂows are also present locally. Unit Tbr is broadly divisible into block-and-ash-ﬂow deposits associated with eruptions at silicic domes and avalanche breccias (Fig. 25A), in turn associated with massive rockfalls from over steepened domes (Fig. 25B). Avalanche breccias are predominantly clast-supported and are composed of pieces of reworked ﬂow banded rhyolite that range in size from ~5 cm to 1 m. Near the base of unit Tbr, avalanche breccia deposits also include clasts of trachydacite. In contrast, the block-andash-ﬂow deposits consist of rhyolite clasts up to 20 cm in size, supported in a matrix dominated by ash and unﬂattened pumice. 40 Ar/39Ar dating of two breccia samples yield ages of 17.68 and 17.34 Ma. The rhyolite breccias with interﬁngered rhyolite ﬂows mark the culmination of felsic volcanism within the Secret Pass Canyon volcanic center. Intrusions such as the trachydacite volcanic neck immediately southeast of Secret Pass Canyon indicate that there was an abrupt transition from felsic to intermediate magmatism at ca. 17.4 Ma, after which time magmatism apparently stopped. The cause for the transition in composition and subsequent cessation in magmatism is unknown and makes for an intriguing question to address in conjunction with studies of plutons exposed to the west across the Colorado River. RETURN to vehicles. Hike west down Secret Pass Canyon wash, moving down section. At the base of the massive upper portion of unit Td, at its contact with unit Tst (bottom of a wet or dry waterfall) (Fig. 26), note a dark vitrophyre. This vitrophyre is the freshest Td we have seen, with pristine phenocrysts of plagioclase, biotite, and clinopyroxene set in a glassy matrix. This vitrophyric base is more consistent with an intrusive contact than a ﬂow base. Beyond this point (down-section to the west) is an extensive, relatively well-exposed section of unit Tst that includes what appear to be lahar deposits. As a ﬁnal point of interest, petroglyphs exist on some rock outcroppings, so please be careful if collecting samples. Once we reach the road, walk right (north) for ~0.8 mi to the vehicles. We will then drive back to Las Vegas.

A

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Figure 23. (A–B) Photos of unit Tvs. (A) Red and white pumiceous sediments that appear to represent airfall tuffs. (B) Ignimbrite unit that locally overlies the pumiceous sediments. Sanidine from this ignimbrite gives an 40Ar/39Ar age of 17.48 Ma (Table 1). Tbr—rhyolite breccias; Tvs—volcaniclastic sediments.

B

unit Tbr

Tuff

unit Tvs

Entrance to Secret Pass Canyon Stratification within breccias?

unit Tbr

unit Tbr

Figure 24. Breccias of unit Tbr at the opening to Secret Pass Canyon. The breccias form massive cliffs, and some appear to contain stratiﬁcation.

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A

Flow banding within some clasts

B

unit Td

unit Tst

Figure 25. (A–B) Close-up photos of the two different types of breccias observed in unit Tbr. (A) Avalanche breccias. These breccias are clast-supported with subangular to rounded clasts of ﬂowbanded rhyolite. Some individual clasts are highlighted by the dashed black line. These deposits likely reﬂect material eroded from rhyolite domes, possibly during growth and collapse. (B) Blockand-ash-ﬂow deposits. These deposits are supported by a matrix of ash and consist of mostly angular clasts of rhyolite. We interpret these deposits as associated with rhyolite dome eruption.

Figure 26. Contact between units Td (trachydacite) and Tst (volcaniclastic sediments and tuffs) as exposed within the Secret Pass Canyon wash. The contact is represented as a baked horizon and has an attitude consistent with the overall regional trend.

Spirit Mountain batholith and Secret Pass Canyon volcanic center ACKNOWLEDGMENTS This ﬁeld trip stems from several years of work in southeastern Nevada–northwestern Arizona and has beneﬁted from the unselﬁsh help of numerous individuals. We would like to thank Jim Faulds for introducing us to the Secret Pass Canyon area and for providing invaluable assistance and insight in the ﬁeld. We thank Warner Cribb for the use of the X-ray ﬂuorescence spectrometer at Middle Tennessee State University. We would also like to thank Heather Bleick, Ben George, Ilmari Haapala, Cliff Hopson, Mark Liggett, Steve Ludington, Rod Metcalf, Jonathan Miller, Zach Miller, Steven Ownby, Tapani Ramo, Rob Smith, Alex Volborth, Bob Weibe, and Joe Wooden for their help and support. We also thank Gene Smith and Ernie Duebendorfer for their work as editors and Keith Howard for his thoughtful review. This work was supported by National Science Foundation grants EAR-0409876 and EAR-0107094. REFERENCES CITED Anderson, R.E., 1971, Thin skin distension in Tertiary rocks of southeastern Nevada: Geological Society of America Bulletin, v. 82, p. 43–58, doi: 10.1130/0016-7606(1971)82[43:TSDITR]2.0.CO;2. Anderson, R.E., 1977, Geologic map of the Boulder City 15-minute Quadrangle, Clark County, Nevada: U.S. Geological Survey Geologic Quadrangle Map, GQ-1395. Anderson, R.E., 1978, Geologic map of the Black Canyon 15-minute Quadrangle, Mohave County, Arizona and Clark County, Nevada: U.S. Geological Survey Geologic Quadrangle Map GQ-1394, scale 1:62,500. Bachl, C.A., Miller, C.F., Miller, J.S., and Faulds, J.E., 2001, Construction of a pluton: Evidence from an exposed cross-section of the Searchlight pluton, Eldorado Mountains, Nevada: Geological Society of America Bulletin, v. 113, p. 1213–1228, doi: 10.1130/0016-7606(2001)113<1213: COAPEF>2.0.CO;2. Bachmann, O., and Bergantz, G.W., 2004, On the origin of crystal-poor rhyolites: Extracted from batholithic crystal mushes: Journal of Petrology, v. 45, no. 8, p. 1565–1582, doi: 10.1093/petrology/egh019. Cas, R.A.F., and Wright, J.V., 1987, Volcanic successions, modern and ancient: A geologic approach to processes, products, and successions: London, Allen and Unwin Publishers, 528 p. Claiborne, L.L., Miller, C.F., Walker, B.A., Wooden, J.L., Mazdab, F.K., and Bea, F., 2006, Tracking magmatic processes through Zr/Hf ratios in rocks and Hf and Ti zoning in zircons: An example from the Spirit Mountain batholith, Nevada: Mineralogical Magazine, v. 70, no. 5, p. 517–543, doi: 10.1180/0026461067050348. Faulds, J.E., 1995, Geologic map of the Mount Davis Quadrangle, Nevada and Arizona: Nevada Bureau of Mines and Geology Map 105, scale 1:24,000, 4 p. text. Faulds, J.E., 1996, Geologic map of the Fire Mountain Quadrangle, Nevada and Arizona: Nevada Bureau of Mines and Geology Map 106, scale 1:24,000, 6 p. text. Faulds, J.E., Geissman, J.W., and Mawer, C.K., 1990, Structural development of a major extensional accommodation zone in the Basin and Range province, northwestern Arizona and southern Nevada, in Wernicke, B.P., ed., Tertiary extensional tectonics near the latitude of Las Vegas: Geological Society of America Memoir 176, p. 37–76. Faulds, J.E., Geissman, J.W., and Shaﬁqullah, M., 1992, Implications of paleomagnetic data on Miocene extension near a major accommodation zone in the Basin and Range province, northwestern Arizona and southern Nevada: Tectonics, v. 11, no. 2, p. 204–227. Faulds, J.E., Feuerbach, D.L., Reagan, M.K., Metcalf, R.V., Gans, P., and Walker, J.D., 1995, The Mt. Perkins block, northwestern Arizona: An exposed cross section of an evolving, preextensional to synextensional magmatic system: Journal of Geophysical Research, v. 100, no. B8, p. 15,249–15,266, doi: 10.1029/95JB01375.

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Nevada: Implications for timescales and physical processes of batholith construction: Journal of Volcanology and Geothermal Research, v. 167, p. 239–262, doi: 10.1016/j.jvolgeores.2006.12.008. Watson, E.B., and Harrison, T.M., 2005, Zircon thermometer reveals minimum melting conditions on earliest Earth: Science, v. 308, no. 5723, p. 841– 844, doi: 10.1126/science.1110873. MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008

The Geological Society of America Field Guide 11 2008

Devonian carbonate platform of eastern Nevada: Facies, surfaces, cycles, sequences, reefs, and cataclysmic Alamo Impact Breccia John E. Warme* Department of Geology and Geological Engineering, Colorado School of Mines, Golden, Colorado 80401, USA Jared R. Morrow* Department of Geological Sciences, San Diego State University, San Diego, California 92182, USA Charles A. Sandberg* U.S. Geological Survey, Box 25046, MS 939, Federal Center, Denver, Colorado 80225, USA

ABSTRACT Devonian limestone and dolostone formations are superbly exposed in numerous mountain ranges of southeastern Nevada. The Devonian is as thick as 1500 m there and reveals continuous exposures of a classic, long-lived, shallow-water carbonate platform. This ﬁeld guide provides excursions to Devonian outcrops easily reached from the settlement of Alamo, Nevada, ~100 mi (~160 km) north of Las Vegas. Emphasis is on carbonate-platform lithostratigraphy, but includes overviews of the conodont biochronology that is crucial for regional and global correlations. Field stops include traverses in several local ranges to study these formations and some of their equivalents, in ascending order: Lower Devonian Sevy Dolostone and cherty argillaceous unit, Lower and Middle Devonian Oxyoke Canyon Sandstone, Middle Devonian Simonson Dolostone and Fox Mountain Formation, Middle and Upper Devonian Guilmette Formation, and Upper Devonian West Range Limestone. Together, these formations are mainly composed of hundreds of partial to complete shallowing-upward Milankovitch-scale cycles and are grouped into sequences bounded by regionally signiﬁcant surfaces. Dolomitization in the Sevy and Simonson appears to be linked to exposure surfaces and related underlying karst intervals. The less-altered Guilmette exhibits characteristic shallowing-upward limestone-to-dolostone cycles that contain typical carbonate-platform fossil- and ichnofossil-assemblages, displays stacked biostromes and bioherms of ﬂourishing stromatoporoids and sparse corals, and is punctuated by channeled quartzose sandstones. The Guilmette also contains a completely exposed ~50-m-thick buildup that is constructed mainly of stromatoporoids, with an exposed and karstiﬁed crest. This buildup exempliﬁes such Devonian structures known from surface and hydrocarbon-bearing subsurface locations worldwide. Of special interest is the stratigraphically anomalous Alamo Breccia that represents the middle member of the Guilmette. This spectacular cataclysmic megabreccia, produced by the Alamo *[email protected]; [email protected]; [email protected] Warme, J.E., Morrow, J.R., and Sandberg, C.A., 2008, Devonian carbonate platform of eastern Nevada: Facies, surfaces, cycles, sequences, reefs, and cataclysmic Alamo Impact Breccia, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 215–247, doi: 10.1130/2008.ﬂd011(10). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION Purpose and Objectives The main purpose of this guide is to present, in outcrop examples, many of the common characteristics of shallow-water carbonate-platform stratigraphic successions. Each stop exhibits carbonate rocks that are typical of the Devonian platform facies of southeastern Nevada and, except for the platform biota that changes through time, are similar in structure to many carbonate platforms that occur elsewhere throughout Phanerozoic and even Proterozoic time. In addition, each stop presents objectives of special stratigraphic interest that sets them apart, such as distinctive cyclic stacking patterns, exposure surfaces, karst intervals, organic buildups, sandstone interbeds, and evidence for cataclysm via the Alamo Breccia. We intend that enough time be dedicated at each stop to promote observations, descriptions, discussions, interpretations, and possibly debates about the rocks. Each stop has been the subject of one or more published papers, or reported in theses, but our approach will be, as much as possible, to examine the formations as a group of unbiased observers and interpreters and interpret the rock properties and their signiﬁcances. BACKGROUND INFORMATION The following brief overviews are presented for the beneﬁt of participants unfamiliar with the ﬁeld trip area and its geographic and geologic setting. Much of this information was compiled by Jared Morrow and presented in an unpublished ﬁeld trip guidebook for participants of the September 2007 meeting of the Subcommission on Devonian Stratigraphy (SDS) of the International Union of Geological Sciences (IUGS). Additional data are excerpted from Morrow and Sandberg (2008). Geographic Setting The ﬁeld trip area lies within a southern portion of the Great Basin (Fig. 1), which is a large part of the Basin and Range physiographic province of western North America. This vast province is characterized by hundreds of named, north-south–trending mountain ranges separated by alluvium-ﬁlled, mainly Tertiary basins. The ranges commonly exhibit tilted and well-exposed Paleozoic stratigraphic sections, as shown in Figure 2, which facilitate study

and can be correlated between ranges. The Great Basin portion includes almost all of Nevada and parts of eastern California and Oregon, southern Idaho, and western Utah. As the name implies, the “Great Basin” is hydrologically deﬁned as the region where streams drain into enclosed basins with no outlets to major rivers, such as the Columbia River system to the north and the Colorado River system to the southeast. Great Basin elevations range from ~600 m in deserts of the southern valleys to 4006 m at the summit of Boundary Peak, the highest point in Nevada. Geologic Overview The oldest rocks in the Great Basin region are Paleoproterozoic, 1.74 Ga metamorphic and intrusive igneous units exposed in southern Nevada. In the east-west corridor from western Utah to central Nevada, this crystalline basement is covered by a westward-thickening prism of Neoproterozoic to Middle Devonian, supratidal to deep-subtidal siliciclastic and carbonate sedimentary rocks that were deposited along the subsiding passive craton margin of what is now western North America (Stewart, 1980). The sedimentary prism reaches a thickness of nearly 9000 m in central Nevada, with Devonian strata making up more than 1800 m of this total (Fig. 3). Regional to global correlation of Devonian rocks, sequences, and events is facilitated by conodont biostratigraphy (Fig. 4), which provides the highest resolution biochronology for this time interval. Figure 5 shows the Middle Devonian to Lower Mississippian formations that occur in the southeastern Nevada area of this ﬁeld guide. They represent sediment accumulation in the eastern and central longitudinal bands of the north-south sedimentary prism. The carbonate-dominated formations represent episodes of a long-lived, shallow-water, carbonate platform that existed until the Late Devonian to Early Mississippian Antler orogeny. Figures 6 and 7 show the relationships between eustasy and Devonian formations that accumulated along the continental margin, including those that are the subject of this Chapter. Figure 7 shows that guidebook stops in the Sevy, Simonson, Fox Mountain, and Guilmette formations (Stops 1–5) lie relatively landward on the platform, whereas the cherty argillaceous unit and Sentinel Mountain, Bay State, and Devils Gate formations (Stop 6) lie more seaward. As outlined below, however, Devonian and later tectonic disturbances, ﬁrst mainly compressive and later mainly extensional, severely fragmented the platform so that paleogeographic reconstructions are difﬁcult and still ongoing.

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Figure 1. Lambert projection digital elevation model shaded-relief image of central Basin and Range province, showing hydrologic boundary of the Great Basin region (dashed white line), outline of Nevada (solid white line), location of selected Nevada cities and towns, and area of ﬁeld trip route shown in Figure 10 (black box). North is toward top of image. Modiﬁed from http://able1.mines.unr.edu.

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Figure 2. View, looking north, of ~700 m of Middle and Upper Devonian formations exposed on a fault block in the West Pahranagat Range. Rocks of this section are almost identical to those exposed on the traverse of Stop 1, along “Downdropped Mountain,” 1 mi (1.6 km) to the west. Devonian carbonate-platform rocks include the Givetian Fox Mountain Formation (Dfm), the Givetian yellow slope-forming member of the Guilmette Formation (Dgysf), the Givetian to Frasnian lower member of the Guilmette Formation (Dgl), the mid-Frasnian Alamo Breccia Member of the Guilmette Formation (Dgab), and the Frasnian to lower Famennian upper member of the Guilmette Formation (Dgu). The Guilmette here contains 70 documented platform carbonate cycles, which are mixed with quartzose sandstone facies in the upper member (Estes-Jackson, 1996).

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Figure 3. Palinspastic map of Great Basin region with isopachs (thickness in thousands of feet) and depositional provinces of Devonian rocks. Position of Lincoln County (L.C.), Nevada, is marked. Modiﬁed from Stewart and Poole (1974).

Beginning in the Middle Devonian and accelerating in the Late Devonian, the effects of eastward-verging tectonic compression crossed the region. Several orogenic episodes were driven by this compression. Most evident are the Antler orogeny (Late Devonian–Mississippian, focused along the Roberts Mountains thrust; Fig. 8), and the Sonoma orogeny (Late Permian–Early Triassic). During the Devonian to Triassic interval, western North America was fringed by a developing subduction zone system. As shown partially in Figures 6 and 7, from east to west across the Great Basin the sedimentary and tectonic provinces at this time included: (1) a shallow-marine, carbonate-dominated platform; (2) a backbulge basin (e.g., the Late Devonian–Early Mississippian Pilot basin); (3) an eastward-migrating forebulge; (4) a foreland basin, which received detrital sediments from both the east and the west; (5) an emergent allochthon; (6) a western, deep-marine basin dominated by hemipelagic sedimentation and maﬁc volcanism; and (7), at the longitude of present-day western Nevada and eastern California, a volcanic arc terrane. Mesozoic sedimentary rocks in the Great Basin are largely of Triassic and Early Jurassic ages, but are not exposed in this ﬁeld guide area. Cretaceous sedimentary rocks are represented

by continental deposits in localized tectonic basins (e.g., the Newark Canyon Formation). By the mid-Jurassic, extensive tectonic compression again spread from west to east across Nevada, culminating in the Cretaceous to early Cenozoic Sevier orogenic belt of easternmost Nevada and western Utah (Fig. 8). During the middle Cenozoic, from ~34 Ma to ~17 Ma, the region was dominated by siliceous volcanism, evidenced by widespread ash-ﬂow tuffs and rhyolites that are preserved on most of the ranges in the ﬁeld trip area. From ~17 Ma to the present, major extensional tectonism, crustal thinning, normal faulting, and maﬁc volcanism characterized the region, resulting in the modern basin and range topography and basaltic ﬂows that can be young enough to follow present drainages. In the Pleistocene, from ~30,000 yr B.P. to ~10,000 yr B.P., large areas of many basins were covered by extensive pluvial lake systems, including Lake Lahontan in western Nevada and Lake Bonneville in western Utah (Stewart, 1980; Hintze, 1988). Wave-cut terraces and ﬁne-grained lacustrine deposits are a common feature in these basins, especially north of the ﬁeld trip area covered in this guide. Also during the Pleistocene, isolated alpine glaciers formed on the highest peaks of the central Great Basin.

Figure 4. Devonian conodont biochronology, scaled to numerical ages of Kaufmann (2006). Correlation of zonation with German Stufen is shown for Late Devonian; Substages are those proposed by Sandberg and Ziegler (1998). Subdivisions of Early rhenana and linguiformis Zones are shown. E—Early; L—Late; semi—semichatovae Subzone interval; MISS.—Mississippian. Scaling of Famennian biozones to numerical ages is approximate and provisional. Modiﬁed from Sandberg et al. (2002), Girard et al. (2005), Kaufmann (2006), and C.A. Sandberg (unpublished).

Figure 5. Time-rock chart of Middle Devonian (Eifelian) to Lower Mississippian (Osagean) stratigraphic units exposed at Alamo Canyon (ALA), and at Bactrian Mountain (BCT and BME) and Silver Canyon, Mount Irish Range (MIR, Stop 3), east and north of Alamo, Nevada. Numerical ages are from Kaufmann (2006). Time values of Alamo Breccia Member, Leatham Member, and Mount Irish buildup are exaggerated for graphic purposes. See Sandberg et al. (1997) for Nevada Events. Modiﬁed from Sandberg et al. (1997).

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Figure 6. Devonian sea-level curve, showing relative paleotectonic and geographic settings of principal stratigraphic units of the western United States. Roman numerals indicate transgressive-regressive (T-R) cycles of Johnson et al. (1985). F.I.—brachiopod- and conodont-based faunal intervals of Johnson et al. (1980); T.S.—transgressive starts in the western United States; C.Z.—conodont Zones. Kobeh, Bartine, and Coils Creek are members of McColley Canyon Formation. From Johnson and Sandberg (1989).

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Figure 7. Northeast-to-southwest, Devonian time-rock transect across central and eastern Nevada, showing carbonate-platform, continental slope, and toe of slope stratigraphic units and relative lateral shifts in platform margin through time. Transgressive-regressive (T-R) cycles Ia through IIf (Johnson et al., 1985, 1991; Johnson and Sandberg, 1989), main intervals of turbidity current and debris-ﬂow deposition (arrows), proto-Antler forebulge initiation (FB), and timing of Alamo Impact Event are indicated. Silty dolostone and siltstone of the yellow slope-forming member (YSF), which forms the basal unit of the Guilmette Formation and Devils Gate Limestone, is a widespread marker lithology distributed throughout western North America (Sandberg et al., 1989, 1997, 2002). The YSF is correlated with ﬁsh-bearing Red Hill beds in north-central Nevada. Four members of Simonson Dolostone are: cxm—coarse crystalline member; lam—lower alternating member; bcm—brown cliff member; and uam—upper alternating member. Other abbreviations: cau—cherty argillaceous unit; Cnyn.—Canyon; Crk.—Creek; Dol.—Dolostone; Fm.— Formation; L.—Lower; Ls.—Limestone; m.—middle; mbr.—member; McMon.—McMonnigal; Mtns.—Mountains; pt.—part; Ss.—Sandstone; t.—tongue; U.—Upper. From Morrow and Sandberg (2008) and based on data from Johnson and Sandberg (1977), Johnson and Murphy (1984), Sandberg et al. (1989, 1997, 2002, 2003), Johnson et al. (1996), and M.A. Murphy (14 September 2007, personal commun.).

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Figure 8. Index map of Great Basin region showing major thrust and strike-slip faults. Sevier thrust system, of Cretaceous to early Cenozoic age, includes individual parts consisting of Nopah, Wheeler Pass, Keystone, Gass Peak, Wah Wah, Muddy Mountains, Blue Mountain, Nebo, Charleston, Willard, and Paris faults. From Stewart (1980).

Devonian carbonate platform of eastern Nevada DEVONIAN GEOLOGIC SUMMARY Rocks studied on this ﬁeld trip represent part of the ~57m.y.-long Devonian geological history of the eastern and central parts of the extensive carbonate platform that extended north-tosouth through Nevada and Utah (Figs. 6, 7). The stops in this guide represent only the central latitudinal segment of a much longer Devonian platform and platform-margin complex that fringed western North America from Alaska and Arctic Canada to Mexico (Ziegler, 1989; Poole et al., 1992). Devonian platformto-basin rocks in western Utah and Nevada record the complex transition from long-term extensional to compressional tectonic modes of the Antler orogeny, which was the ﬁrst in a series of late Paleozoic to Mesozoic orogenic belts that developed over the subduction zone fringing the western continental margin of North America. In the middle- to outer-platform settings, Devonian strata reach a thickness of over 1800 m (Fig. 3). During the mid-1970s to early 1990s, a series of landmark papers provided detailed stratigraphic, biostratigraphic, lithofacies, isopach, structural, and coastal onlap data on the Devonian continental margin in the western United States, including the eustatic sea-level curve of Figure 6 and multiple paleogeographic and paleotectonic time-slice maps constrained by highresolution conodont biochronology. These important papers include, among others: Poole (1974), Stewart and Poole (1974), Johnson and Sandberg (1977, 1989), Matti and McKee (1977), Murphy (1977), Poole et al. (1977, 1992), Sandberg and Poole (1977), Johnson and Pendergast (1981), Gutschick and Sandberg (1983), Kendall et al. (1983), Johnson and Murphy (1984), Murphy et al. (1984), Johnson et al. (1985, 1986, 1989, 1991), Stevens (1986), Sandberg et al. (1988, 1989), Goebel (1991), Johnson and Bird (1991), D.M. Miller et al. (1991), and E.L. Miller et al. (1992). More recent studies of the Devonian platform and platformto-basin transition in Utah and Nevada have focused on reﬁning aspects of: (1) tectonic and structural history (e.g., Oldow et al., 1989; Giles, 1994; Giles and Dickinson, 1995; Crafford and Grauch, 2002; Grauch et al., 2003; Sandberg et al., 2003); (2) depositional models and facies geometry (e.g., Elrick, 1995, 1996; LaMaskin and Elrick, 1997; Cook and Corboy, 2004); (3) conodont-based event stratigraphy and eustasy (e.g., Johnson et al., 1996; Sandberg et al., 2002, 2003; Morrow and Sandberg, 2003, 2008); (4) stratigraphic and structural development in relation to synsedimentary exhalative gold mineralization in the Carlin gold trend, northern Nevada (e.g., Emsbo et al., 1999, 2006; Hofstra and Cline, 2000; Crafford and Grauch, 2002); and (5) relation to the early Late Devonian Alamo Impact Event, south-central Nevada (e.g., Warme and Sandberg, 1995, 1996; Sandberg et al., 1997, 2002, 2003, 2005, 2006; Warme and Kuehner, 1998; Warme, 2004; Morrow and Sandberg, 2005, 2006; Morrow et al., 2005; Pinto, 2006; Warme and Pinto, 2006; Pinto and Warme, 2008). Warme and Pinto (2006) and Pinto and Warme (2008) presented a genetic classiﬁcation for different facies of the Alamo

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Breccia, placing known positions of the Breccia into impact “Realms” with respect to the target zone. As shown in Figure 9, Stops 1–5 in this chapter fall in the Ring Realm, and Stop 6 represents the only known locality in the Rim Realm, closest to the as yet unidentiﬁed central target zone. The other realms are outside the area of this chapter, but have been described and interpreted in references cited for the Alamo Impact Event. Sedimentation Patterns In western Utah and Nevada, the position of the Devonian carbonate platform and the sedimentary facies deposited in platform-to-basin settings were inﬂuenced by both eustasy and tectonics. Since the early comprehensive work of Roberts et al. (1958), Paleozoic sedimentary strata in western Utah and east-central Nevada have been generally regarded to document, from eastto-west, shallow-to-deep marine carbonate-platform, platformmargin, slope, and basin depositional settings (Figs. 5 and 7). In general, Devonian platform deposits are dominated by supratidal, intertidal, and shallow-subtidal carbonate rocks such as biolaminated dolostone, bioturbated lime mudstone and wackestone, bioclastic wackestone and packstone, stromatoporoid-dominated lime mud-rich biostromes, and intraformational conglomerate (Elrick, 1996; Poole et al., 1992; Cook et al., 1983; Cook and Corboy, 2004). During eustatic lowstands, craton-derived quartz sand was deposited in channels and basinward-prograding clastic wedges across large areas of the platform (Fig. 5). Devonian outer-platform to platform-margin rocks consist of shallow- to deep-subtidal, bioturbated, bioclastic wackestone, packstone, and grainstone. During parts of the Early and Middle Devonian when a rimmed platform margin developed (Elrick, 1996; Cook and Corboy, 2004), deposits included coral- and crinoid-dominated biostromes, bioherms, and mudmounds. Slope and basin areas accumulated bioclastic and sandy packstones and grainstones deposited within submarine debris-ﬂow and turbidite-fan systems, intraformational slumps with ﬂat-clast conglomerates, and rhythmically deposited argillaceous carbonate and ﬁne- to medium-grained siliciclastic units. The most distal deposits include rhythmically bedded radiolarian chert, very ﬁne- to ﬁne-grained siliciclastic units, and minor greenstones. The Early and Middle Devonian carbonate platform-to-basin transition, which can be in part characterized using classic carbonate platform models (e.g., Wilson, 1975; Read, 1982), was characterized by several morphologies including homoclinal ramps, distally steepened ramps, rimmed platform margins with intra-shelf basins, and rimmed platform-margins ﬂanked by landward shallow-subtidal platforms and seaward slope-aprons (Cook et al., 1983; Kendall et al., 1983; Johnson and Murphy, 1984; Schalla and Benedetto, 1991; Elrick, 1996; Cook and Corboy, 2004). As discussed next, however, throughout the latter half of the Devonian, proto-Antler and Antler tectonism exerted a strong to dominant control on the position and geometry of the carbonate-platform marginto-basin settings.

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Figure 9. Alamo Breccia locality map showing genetic Breccia Realms, Nevada and western Utah. Lateral Breccia Zones 1, 2, and 3 of earlier publications (e.g., Warme and Sandberg, 1995, 1996; Warme and Kuehner, 1998) are now equated with the Rim, Ring, and Runup Realms (Warme and Pinto, 2006; Pinto and Warme, 2008), respectively. Black diamonds, offshore, deep-water Alamo channel localities; white circles, localities at Tempiute Mountain interpreted to be on or within the Alamo crater rim; open triangles, carbonate-platform localities with well-developed Alamo Breccia including potentially all Units A–D; black circles, carbonate-platform localities with thin Alamo Breccia; black triangle, locality with seismically disturbed zone stratigraphically equivalent to Alamo Breccia; white diamonds, distal, middle- to inner-platform Alamo channel deposits; black arrows, paleocurrent directions determined from clast imbrication; black squares, selected towns. Field trip Stops 1–6 are indicated by numbers. Modiﬁed from Pinto and Warme (2008).

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Tectonic Models Neoproterozoic to early Paleozoic time was marked by a complex pattern of tectonic extension and rifting along western North America as the proto-Paciﬁc basin opened following the breakup of Rodinia at ~850–650 Ma (Stewart, 1972; Stewart and Suczek, 1977; Poole et al., 1992). Outer-platform and platformmargin basins, which probably formed by reactivation of structures inherited from Mesoproterozoic to Neoproterozoic rifting of underlying crystalline continental basement (Stewart, 1972; Stewart and Poole, 1974), strongly inﬂuenced sedimentation patterns in Nevada during the Middle Cambrian to Middle Devonian (Johnson and Potter, 1975; Matti and McKee, 1977; Johnson and Murphy, 1984; Miller et al., 1991; Poole et al., 1992). For the Early and Middle Devonian interval especially, the distribution of facies patterns, rock isopachs, strontium and lead isotope isopleths,

basement gravity signatures, and synsedimentary exhalative gold and barite deposits all suggest that the platform margin was characterized by a complex series of restricted, active, fault-bounded sub-basins, which determined the type, extent, and thickness of sedimentary units (Grauch, 1998; Emsbo et al., 1999, 2006; Hofstra and Cline, 2000; Crafford and Grauch, 2002; Grauch et al., 2003; Emsbo and Morrow, 2005; Morrow and Sandberg, 2008). By the Ordovician (Ross, 1977) or Silurian (Poole et al., 1977), a volcanic island-arc system developed over subducted oceanic crust west of the North American continent. Localized extensional or transtensional tectonics within a postulated inner-arc basin located between the volcanic arc and the continent may have further promoted the formation of fault-bounded platform-margin sub-basins prior to late Middle to early Late Devonian compression and transpression associated with the approaching Antler orogen (Poole et al., 1977; Eisbacher, 1983; Crafford and Grauch, 2002).

Devonian carbonate platform of eastern Nevada Unraveling of the complicated Devonian tectonic processes and sedimentary responses in the Great Basin area has been facilitated by use of conodont biostratigraphy (e.g., Figs. 4, 6, and 7). The earliest direct stratigraphic evidence for the switch to a convergent tectonic mode is in central Nevada, where uplift and erosion was associated with development of the initial, shallow-marine to emergent, proto-Antler forebulge during the latest Givetian to early Frasnian disparilis, falsiovalis, and transitans conodont Zones (Figs. 6 and 7; Sandberg et al., 2003). Subsequent Late Devonian depositional settings and facies patterns on the outer platform and platform margin were dominated by tectonic effects of the converging Antler orogenic belt, which formed an eastward-migrating system composed, from west to east, of an allochthon, a foreland basin, a forebulge, and a backbulge (Pilot) basin that developed across the carbonate-dominated platform to the east (Poole and Sandberg, 1977; Goebel, 1991; Giles, 1994; Giles and Dickinson, 1995). By the late middle Famennian Early postera Zone (Fig. 6), the carbonate platform environment was terminated by widespread uplift driven by the continued eastward migration and expansion of the Antler orogen. Erosion off the leading forebulge formed a regional unconformity that interrupted or removed the latest Famennian depositional record of the platform (Sandberg et al., 1989, 2003; Poole and Sandberg, 1991). The magnitude of this extensive unconformity was ampliﬁed by a major eustatic sea-level fall that began during the late Famennian Middle praesulcata Zone and persisted into the Early Mississippian (Sandberg et al., 1989; 2002). At the site of the former carbonateplatform margin, the overlying Mississippian (Kinderhookian to Chesterian) foreland basin fan and overlap assemblage rocks include compositionally immature, syntectonic, siliciclastic units that were derived in large part from the Antler allochthon to the west (Johnson and Pendergast, 1981; Poole and Sandberg, 1991). Final convergence of the Antler orogen with western North America during the Early Mississippian thrust Devonian and underlying lower Paleozoic basin and slope rocks as much as 145 km eastward over coeval shelf-margin and outer-shelf rocks, forming the Roberts Mountains thrust system (Fig. 8; Stewart, 1980; Johnson and Pendergast, 1981; Poole et al., 1992). Recent models of the Roberts Mountains thrust system characterize it as a complex zone of intercalated, folded, thrust, and imbricated upper Precambrian to middle Paleozoic structural-stratigraphic units (e.g., Theodore et al., 1998; Noble and Finney, 1999; Crafford and Grauch, 2002). These relationships greatly complicate efforts to precisely reconstruct the paleogeography of the depositional systems that operated at this time. CARBONATE PLATFORM FACIES Hundreds of books, monographs, symposium volumes, and journal articles have been devoted to the sedimentology, stratal geometries, biota, petrology, diagenesis, resource potential, and other aspects of modern and ancient shallow-water carbonate platforms. The following brief overview is intended to highlight some

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of the characteristics of carbonate platforms that are exhibited at stops described in this guide. Only a few references are cited herein, but the interested reader is encouraged to consult them and choose from the abundant cited references that they provide. Background Many early reports on carbonate rocks were exhaustive petrographic studies of the rock and fossil particles that compose limestones and dolostones of different ages, and were published in German, French, Italian, Russian, and other languages, as well as in English. Much of the non-English material was ignored or unappreciated by American geologists. By the 1970s, enough understanding was achieved to synthesize the accumulated ﬁeld, subsurface, and petrographic studies into facies models that could help predict the distribution of rock types in carbonate depositional systems. Research on existing shallow-water carbonate environments, such as the Bahama Banks, the Yucatan-Belize shelf, and the Great Barrier Reef, together with regional work on classic ancient localities such as the Devonian reef complex of northwest Australia, the Permian reef complex of west Texas, and the Triassic exhumed atolls of the Italian Dolomite Alps, all helped establish the “carbonate platform” paradigm as the basic framework for the accumulation of most marine shallow-water carbonate rocks. Two classic volumes are fundamental resources for carbonate workers. Wilson (1975) comprehensively synthesized results from studies worldwide and showed how carbonate platforms functioned as similar sediment-accumulation systems through time. He demonstrated how platforms of different ages had similar three-dimensional forms but were generated by different organisms that adapted and evolved over time. In addition to corals and algae that dominate modern platform seaward margins, older margins and rims were constructed by various carbonateﬁxing organisms, such as bacteria, sponges, bryozoans, brachiopods, tube-building worms, and bivalve and gastropod mollusks. From the stromatolitic environments in the Proterozoic to the coral-algal (“coralgal”) reefs of today, platform carbonates accumulated laterally and stacked vertically in response to relative sea-level changes that controlled the availability of accommodation space for the in situ generation of new sediment. The compendium edited by Scholle et al. (1983) brought together the knowledge of varied carbonate depositional environments, from continental and lacustrine to pelagic deep-marine, and offered several chapters on the supratidal to subtidal environmental bands that occur across carbonate platforms. Many other volumes dedicated speciﬁcally to carbonate platforms followed (e.g., Crevello et al., 1989; Tucker et al., 1990; Simo et al., 1993). Although some carbonate-dominated shorelines lead seaward down gently inclined ramps, important facies faunas were those that built and maintained a seaward shallow-platform rim into the active wave zone. A very narrow band of reefoid facies, along wave-inﬂuenced seaward platform margins, separates the expansive and relatively quiet lagoonal environments landward of the rims from the steep, debris-covered aprons seaward of them.

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Carbonate accumulations commonly become cemented directly on the seaﬂoor, most notably strengthening the critical narrow band of reefoid and related facies along seaward margins. Bored hardgrounds, sharply eroded transgressive surfaces, short-term disconformities, and other evidence show that seaﬂoor cementation and early burial cementation and diagenesis may completely lithify the sediments of each short-term sediment accumulation, or cycle, prior to initiation of the following one. The Carbonate Platform Signature: Cyclostratigraphy The basic building block of shallow-water carbonate, evaporite, and siliciclastic sedimentary deposits is generally agreed to be the “shallowing-upward cycle,” which has been described from worldwide examples in rocks of Proterozoic to Recent ages (e.g., Ginsburg, 1975). Carbonate platforms commonly exhibit obvious stacked meter-scale cycles whose analysis has given rise to the discipline of “cyclostratigraphy” (e.g., Elrick, 1995). A cycle is generated when ﬂooding creates accommodation space that is invaded by carbonate-producing organisms, which establish the “carbonate-generating factory” in situ across the platform. A fully preserved idealized cycle is a vertically stacked package of genetically related beds of subtidal, intertidal, and supratidal facies, bounded by exposure surfaces that form between cycles (e.g., Hardie and Shinn, 1986; chapters in Loucks and Sarg, 1993). Of course, each cycle has an intertidal to supratidal feather edge that limits the internal sequence of facies, and every platform at any moment has a particular mosaic of shoals, channels, and clusters of carbonate-producing and sediment-trapping organisms. Thus, cycles are expected to show internal lateral variation. Nevertheless, numerous examples have been documented where an individual distinctive cycle, or bundle of cycles, was traced laterally for many kilometers without signiﬁcant change in their interval character. We will see such examples in ﬁeld traverses through the Guilmette Formation during this excursion. Debate about carbonate rock cycles, whether they are driven mainly by external (allocyclic) forces such as climate periodicity and global sea-level response, or internal (autocyclic) processes such as local diastrophism and expected sediment dispersal and accumulation patterns, has prompted numerous studies of modern and ancient platforms coincident with a range of theoretical models. Increased computing power has allowed development of increasingly sophisticated models. A recent example are the models of Burgess (2006), whose paper also contains a comprehensive review and reference list for the history of the extrinsic/ intrinsic debate, which has not yet been fully resolved. With the advent of “sequence stratigraphy” (e.g., Wilgus et al., 1988), researchers tested stratal patterns for hierarchical arrangements that indicated response to relative sea-level changes on various time scales, most notably within the astronomically driven Milankovitch time band in which climate change oscillated at four or more different astronomically ﬁxed intervals, each a few tens of thousands to hundreds of thousands of years in duration (cf., papers in Arthur and Garrison, 1986). The oscillations are not syn-

chronized, so that they may cancel or reinforce each other through time. However, the concept of sedimentary cycles, driven by relative sea-level changes, regardless of cause, is widely accepted. In sequence-stratigraphic terms, a simple system of oscillating sea level together with continuous subsidence could account for the generation and preservation of such cycles (Sarg, 1988). Packages of cycles were documented whose facies prograded, retrograded, and stacked, creating successions of shallow-water platform carbonates that in some cases were hundreds or even thousands of meters in thickness, shown or presumed to be preserved behind some form of subsiding platform rim. Cycles of shorter duration within the Milankovitch band, the ~20,000 yr or ~40,000 yr oscillations, may progress faster than carbonate sediment can accumulate. Some cycles exhibit evidence for transgressive drowning that deepened the platform to below wave base, or regressive shallowing that exposed the platform top to karstiﬁcation and pedogenesis. Because the sediment cycles are commonly only a few meters thick, they may represent a compressed record of the maximum sea levels that occurred during their genesis. In such cases, sea level rose at a faster pace than sediment could accumulate, then fell to meet the surface of aggrading sediments before the maximum accommodation space was ﬁlled. Thus, the exact magnitude of maximum relative sea-level change over a platform during the life of a cycle cannot be fully known, and is a product of local subsidence, eustatic change that may have occurred, and sediment accumulation rate. Factors that further inﬂuence sediment accumulation in any cycle include the somewhat unpredictable mosaic carbonate production and accumulation across any given platform, and the results of rare events such as storms. However, together with the preserved cycle thickness, the magnitude of bathymetric change can be estimated using lithofacies and biofacies depth indicators that may be captured and compressed within a cycle. Within any vertical succession, several similar platform cycles may be bundled, separated from other bundles by a surface that represents extended exposure or an interval that indicates prolonged drowning. Such bundles have been interpreted, not without controversy, to represent shorter-term Milankovitch cycles (tens of thousands of years) captured within longer-term ones (hundreds of thousands of years). However, platform cycles commonly appear to be sorted into bundles of similar internal character. More signiﬁcant regional surfaces of exposure or drowning across platforms may serve to vertically partition bundles of cycles into longer-term sequences. Thus, platform carbonate rocks commonly exhibit a hierarchy of shallowing-upward cycles, separated by a hierarchy of surfaces. The life of a platform may cease by a pronounced relative drop of sea level, exposure, and karst development, or by marine drowning, from which the shallow-water platform environment, with its cycle-generating mechanism, never recovers. Examples in the ﬁeld trip area include the terminal exposure and karst formation at the top of the Simonson Dolostone, and the deepening and extinction of the shallow platform at the top of the Guilmette Formation; both examples are exhibited on the traverse of Stop 1. In Nevada, Devonian carbonate-platform cycles have been studied in detail, arranged into sequences, and analyzed for the

Devonian carbonate platform of eastern Nevada transgressive-regressive sea-level history that they may reveal. North of the ﬁeld trip area, bed-by-bed analyses include those of Elrick (1995, 1996) for Lower and Middle Devonian formations, and LaMaskin and Elrick (1997) for the Guilmette Formation. Within the ﬁeld trip area, Estes-Jackson (1996) provided similar descriptive detail and interpretation of cycles in the Guilmette exposed in the fault block (Fig. 2) adjacent to the Hancock Summit West location of Stop 1. Chamberlain and Warme (1996) and Chamberlain (1999) developed the sequence analysis summarized in the composite stratigraphic column of Figure 11 and the accompanying Table 1. FIELD TRIP AREA Figure 10 shows the ﬁeld trip area and location of stops we have chosen to include in this ﬁeld guide. Numerous other localities, in ranges within and beyond the trip area, also contain excellent exposures of Devonian formations. Prior ﬁeld trip guides that cover some or all of the area are those of Sandberg et al. (1997), which was extensively drawn upon for the present guide, and Gillespie and Foster (2004), which also contains seven reprinted papers on the Alamo Breccia. This guide is the ﬁrst to focus mainly on the Devonian platform beds and their interpretation. Stop Locations The stop locations described herein may be visited independently, in any order. However, Stop 1, Hancock Summit West, contains a thick, continuous section of Devonian formations that are accessible, well exposed, and typical of the ﬁeld trip area. Hence, Stop 1 is most useful as an initial section for comparisons with stratigraphic units at other stops and beyond. Starting Point The location of each stop is described as the direction and distance from a central point in the ﬁeld trip area: a wayside rest near Crystal Springs, at the junction of Nevada State Highways 375 and 318, at the southern edge of the small settlement of Hiko (Fig. 10). Global Positioning System (GPS) coordinates at the Crystal Springs rest area are: lat 37°13′57.20″ N., long 115°24′56.75″ W., Hiko 7.5′ quadrangle. DEVONIAN FORMATIONS Devonian formations to be studied are the Lower Devonian Sevy Dolostone, Lower to Middle Devonian cherty argillaceous unit and Oxyoke Canyon Sandstone, Middle Devonian Simonson Dolostone and Fox Mountain Formation, Middle to Upper Devonian Guilmette Formation, and Upper Devonian West Range Limestone. See Figures 6 and 7 and cited papers for precise ages, lateral equivalents, and interrelationships of these formations. Figure 11 is a composite column of formations, nearly 5000 ft (~1500 m) thick, from the upper part of the Sevy

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Dolostone to the post-platform Devonian to Mississippian Pilot Shale and Lower Mississippian Joana Limestone. The column represents an attempt to place the Devonian formations in the ﬁeld trip area into a sequence-stratigraphic framework. It was constructed by Alan Chamberlain (Chamberlain and Warme, 1996; Chamberlain, 1999) on the basis of outcrops in the central Timpahute Range, close to Stop 3. The depicted sequence differs in detail from sequences exposed at other stops. For example, at Stop 3 the upper Guilmette Formation, shown in the column of Figure 11, contains a thick carbonate buildup (Dgb3), directly over the Alamo Breccia (Dgb2), and only sparse, thin quartz sandstone beds. In contrast, the upper Guilmette at Stop 1 contains thick sandstone intervals and lacks a buildup, whereas at Stop 4 it contains both small buildups and sandstones over the Breccia. The central column of Figure 11 shows basic rock types, partitioned vertically into sequences. The sequences were deﬁned in two ways. The ﬁrst method was to identify and describe individual shallowing-upward cycles in as much detail as possible from available outcrops. This process resulted in bundles of two to as many as 29 cycles of generally similar character. These cycles are listed on Table 1 and shown as excursions on the graphed line to the right of the column. The second method was to create a companion log of the outcrops using a hand-held gamma-ray scintillometer, shown as the graphed line to the left of the column. The two methods were combined to deﬁne sequences that contained cycles of similar lithologic character and gamma-ray signature, separated from other sequences by distinctive but commonly subtle surfaces that signify exposures, signiﬁcant marine transgressions, or other environmental shifts. The graphed line to the right of the column was interpreted as a rough proxy for relative sea-level changes across the platform (for details see Chamberlain and Warme, 1996; Chamberlain, 1999). Characteristics of these sequences are discussed in the text for the stops. Note, however, that many of the sequence boundaries tend to match the boundaries of named formations and members, from the upper Sevy Dolostone (Dse3) at the base to the carbonate buildup over the Alamo Breccia (Dgb3) at the top. The overlying upper segment of the informal upper member of the Guilmette Formation was partitioned into ﬁve thick sequences (Fig. 11), some of which could be useful for further subdivision. This example shows an important attribute of sequence stratigraphy, whereby rock bodies are bounded by genetically signiﬁcant surfaces that mark bathymetric shifts or other environmental events. Boundaries of many members, lenses, and tongues, which originally may have been loosely characterized by terms such as “transitional” or “ﬁrst occurrence” of a given lithology, can be reﬁned, redeﬁned, or clariﬁed if such surfaces are present and recognized. Guilmette Formation All stops described in this guide, except Stop 2, include the Guilmette Formation, which is the least dolomitized and

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115°

116°

38°

0

50 mi.

0

80 km

6 3

2 4 Crystal Springs

1

OR

Alamo

5

ID

37° NEVADA White Pine Co.

UT

Nye Co.

CA

Lincoln Co. N 0

100 km

Detailed map

Clark Co.

AZ

Figure 10. Index and highway map of southeastern Nevada, showing locations of ﬁeld trip Stops 1–6 (numbered stars) and Crystal Springs, the starting point for ﬁeld trip road logs. Modiﬁed from: http://www.nevadadot.com/traveler/maps/StateMaps.

Devonian carbonate platform of eastern Nevada most lithologically variable formation because it contains the Alamo Breccia Member. The Guilmette and its members at Stop 1 are shown in Figures 11–14. The following review is a history of studies and proposed lithostratigraphic subdivisions of this formation. Initially, Reso (1963) divided the Guilmette into informal lower and upper members. The lower member terminated at the top of a thick, stromatoporoid-rich breccia that was documented at Hancock Summit West (Stop 1), along strike in Guilmette exposures on the west side of the West Pahranagat Range, and also near Mount Irish (Stop 3). At Mount Irish, a similar breccia lies directly under a ~50-m-thick stromatoporoid-rich mound (“Reso’s Reef”). Consequently, Reso interpreted that breccia as mound talus, as did Dunn (1979), whose main objective was the study of the mound proper. Because the thick breccia at Hancock Summit West contains abundant stromatoporoids, it was probably also regarded as reef debris, although no large reef is exposed there. At both localities the thick breccia is the Alamo Breccia Member. Estes-Jackson (1996) described and interpreted a Guilmette stratigraphic section that is exposed on the fault block adjoining Stop 1, one mile east of Hancock Summit West (see Fig. 2). The section is almost identical to that traversed at Stop 1. She identiﬁed 22 shallowing-upward cycles in Reso’s lower member and 47 cycles in his upper member. She believed that the thick breccia at the top of the lower member was a normal, but perhaps deeper, platform facies and not a cataclysmic bed; however, it too is the Alamo Breccia Member. She calculated that the cycles fell within the lower part of the Milankovitch band, each of less than 100,000 years duration. Estes-Jackson (1996) documented the sandstone facies of the upper Guilmette, which is well represented at Stop 1 and in many other ranges of eastern Nevada, but is largely absent from the nearby central Timpahute Range where the stratigraphic column of Figure 11 was generated. Kuehner (1997) subdivided Reso’s (1963) lower member of the Guilmette Formation into three units, which are exposed in several ranges: a basal, yellow-weathering “slope forming interval,” a “ledge forming interval,” and the then newly designated Alamo Breccia Member. He retained Reso’s upper member for the balance of the Guilmette. Sandberg et al. (1997) partitioned Reso’s (1963) lower member of the Guilmette into three members (Fig. 14). Their “yellow slope-forming” and “carbonate platform facies” members are similar to Kuehner’s two lower “intervals.” They formalized the highest interval as the type Alamo Breccia Member of the Guilmette Formation. However, elsewhere the widespread cataclysmic Alamo Breccia rests not on the Guilmette, but on an erosive surface cut into Middle Devonian formations. At Stop 6 (Fig. 10) the Breccia takes the form of Units interpreted as fallback and resurge breccias associated with the Rim Realm (Pinto and Warme, 2008), and offshore to the west as deep-water channels of the Runout/Resurge Realm (Fig. 9) that were described at several localities by Morrow et al. (2005) and Sandberg et al. (2005, 2006).

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Sandberg et al. (1997) also described the lower part of Reso’s upper member, shown as the “slope facies member” on Figure 14. The four Guilmette members shown on Figure 14 total ~220 m in thickness. However, at Stop 1 the entire Guilmette totals ~660 m. Our traverse will include the upper ~440 m, which is highly heterolithic but contains classic facies that are signatures of Devonian carbonate platforms. This interval has not been described in detail on “Downdropped Mountain,” but a description in the adjacent fault block is available (Estes-Jackson (1996). Chamberlain and Warme (1996) and Chamberlain (1999) provided the composite stratigraphic column of Figure 11. The Devonian platform formations were divided into sequences, deﬁned by bundles of similar cycles or trends in cycles. These are brieﬂy described in Table 1. In general, the interval from the base of the Guilmette to the top of the Alamo Breccia, Reso’s lower member, is similar across several ranges in southeastern Nevada, but the remainder of the Guilmette, Reso’s upper member, exhibits more variation between ranges as well as within the member. Alamo Breccia Member of Guilmette Formation Various facies of the Alamo Breccia Member of the Guilmette Formation are traversed at Stop 1 and Stops 3–6 and are described in this guide. The Breccia has been treated in detail in other guides and papers, so it is described herein with only enough detail to satisfy our purposes of placing it in stratigraphic context, showing its characteristics, and conﬁrming its genesis as a cataclysmic deposit created by a bolide (i.e., a large craterforming projectile such as an asteroid or comet). Figure 9 shows the general distribution of Alamo Breccia outcrops, grouped into genetic Realms. Stop 1 and Stops 3–5 lie within the Ring Realm, where the Breccia is expressed as four sequential units of distinctive facies, which are labeled as parts of the Alamo Breccia Member on Figure 14. These units were described in detail, initially by Warme and Sandberg (1995, 1996), most recently by Pinto and Warme (2008), and in several intervening reports. Two of the units occur together in the Breccia. The lowest one, termed Unit D in past reports, is a detachment monomict breccia that formed between the undamaged carbonate platform beds below and Unit C megaclasts composed of displaced but intact cyclic platform beds above. The two remaining Alamo Breccia units are polymict breccias of chaotically bedded Unit B, which may extend to the base of the Breccia, as shown in Figure 14, and sorted and graded beds of Unit A that everywhere top the Breccia. For more in-depth study and understanding of the Breccia, we suggest that this ﬁeld guide be augmented by the following easily obtained references, copies of which are provided to participants of the ﬁeld trip for which this guide was prepared: Warme and Sandberg (1996), Sandberg et al. (1997, 2005), Warme and Kuehner (1998), Morrow and Sandberg (2001), Morrow et al. (2001, 2005), Warme et al. (2002), Warme (2004), Warme and Pinto (2006), and Pinto and Warme (2008).

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A

Figure 11 (on this and following page). (A) Composite stratigraphic column of Devonian section near Silver Canyon (Stop 3), showing sequences, surface gamma-ray log, relative sea-level curve, and sequence-boundary features. (B) Legend for sequence symbols, boundary features, and lithologic symbols used in Figure 11A. From Chamberlain and Warme (1996) and Chamberlain (1999). See Table 1 for sequence thicknesses, numbers of cycles, and signiﬁcant features.

Devonian carbonate platform of eastern Nevada

B

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Figure 11 (continued).

TABLE 1. THICKNESSES, NUMBERS OF CYCLES, AND SIGNIFICANT FEATURES OF DEVONIAN SEQUENCES NEAR SILVER CANYON (STOP 3), AS SHOWN IN THE STRATIGRAPHIC COLUMN OF FIGURE 11 Sequence abbreviation MDp2 MDp1 Dwr Dgg Dgf Dge Dgd Dgc Dgb3 Dgb2 Dgb1 Dga2

Thickness ft (m) 115 (35) 130 (39) 153 (46) 567 (172) 268 (81) 235 (71) 406 (123) 188 (57) 97 (29) 179 (54) 26 (8) 145 (44)

No. of cycles 2 2 4 29 16 17 23 6 2 1 2 8

Dga1

250 (76)

12

Dgys

182 (55)

10

Dgfm Dsiualt

135 (41) 285 (86)

4 12

Dsibc Dsilalt Dsicxln

85 (26) 265 (80) 225 (68)

4 12 4

Dox2 Dox1

95 (29) 100 (30)

2 4

Dse3

240+ (73+)

12+

Significant features; weathering profile Silicified stromatolites and laminated black chert; slope Silty limestone capped with fossil bone-bearing sandstone; slope Silty, burrowed limestone; partly covered slopes Carbonate cycles capped by thick (>3 m) quartz sandstone beds Slightly deeper cycles and more limestone than in adjacent sequences Carbonate cycles capped by thin (<3 m) quartz sandstone beds Amphipora dolopackstone; dark-gray ledges and cliffs Silty limestone with abundant gastropods and burrows; slope Stromatoporoid and coral reef facies; light-gray cliffs Graded bed of carbonate breccia, open-marine fauna; brown-gray cliffs Abundant corals, stromatoporoids, and Amphipora; limestone cliffs Shallowing-upward cycles that successively deepen upward, predominately limestone, open-marine fauna; ledges and slope Shallowing-upward cycles that successively deepen upward, predominately dolostone, open-marine fauna; ledges and slope Yellow, silty dolostone, stromatolites, and cycles capped by thin beds of very fine-grained quartz sandstone, ostracodes; slope Open-shelf fauna, brachiopod Stringocephalus; resistant cliffs Shallowing-upward cycles that successively deepen upward giving an alternating dark and light band appearance, karst breccia; ledges Open-shelf fauna with corals and stromatoporoids; dark brown-gray cliff Alternating intertidal-supratidal or dark and light bands; ledges Coarsely crystalline dolostone capped by karst surface; light-gray to light graybrown cliffs Quartz sandstone with hummocky cross-bedding at base; ledge Burrowed, silty dolostone with flat-pebble conglomerate at base; light-brown slope Light-gray, fine-grained, laminated dolostone; slopes, base concealed

TOTAL 4370+ (1324+) 188+ Note: Modified from Chamberlain and Warme (1996) and Chamberlain (1999).

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Mj

Tv

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p MD

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b

Dga

Dgs

p

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Dgl sf Dgy

375

ay hw ig lH ria

Ex tra ter res t

Acc es s road

Dfm t ual Dsi bc

Dsi

Figure 12. Oblique Google Earth aerial view (eye altitude: 2.2 km) to the southeast of “Downdropped Mountain,” Stop 1 at Hancock Summit West, showing access road, traverse route, and stratigraphic units including brown cliff (Dsibc) and upper alternating (Dsiualt) members of Simonson Dolostone, Fox Mountain Formation (Dfm) undivided, and yellow sloping-forming member (Dgysf), lower member (Dgl), type Alamo Breccia Member (Dgab), and upper member (Dgu) of Guilmette Formation. Dgcp and Dgs denote carbonate platform and slope facies, respectively, of Guilmette Formation shown in the stratigraphic column (Fig. 14). Other units: MDp, Pilot Shale; Mj—Joana Limestone; Tv—Tertiary volcanic rocks. Image modiﬁed from http://earth. google.com/.

Stop 1 Start traverse

Figure 13. West face of Hancock Summit West (Stop 1), with stratigraphic sequence including the upper part of Simonson Dolostone (Dsi), Fox Mountain Formation (Dfm), and yellow slopeforming member (YSF), lower member (Dgl), type Alamo Breccia Member (lower Units D–C and upper Units B– A), and upper member (Dgu) of Guilmette Formation. The Alamo Breccia Member is ~55 m thick. The Stop 1 traverse ascends the ridge along the skyline from right to left (Fig. 12). Stop 1 access road is visible in foreground.

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Figure 14. Stratigraphic section, facies, conodont biostratigraphy, and conodont biofacies at Hancock Summit West (Stop 1; Figs. 12 and 13) of upper member of Fox Mountain Formation and lower part of Guilmette Formation, showing type section of Alamo Breccia Member. Shows position of 24 conodont samples used to constrain biostratigraphic age. Conodont biofacies abbreviations: Icr.—icriodid; Pol.-icr.—polygnathidicriodid; No c.—no conodonts. From Sandberg et al. (1997).

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FIELD TRIP STOPS Stop 1. Hancock Summit West: Middle and Upper Devonian Formations, and Type Section of the Alamo Breccia Location Stop 1 is at the informally named “Downdropped Mountain” or “Down Dropped Block” (Fig. 12), which is a prominent ridge south of, and parallel to, Highway 375, from 12–14 mi (19–22 km) west of the starting point at the Crystal Springs rest area. The entrance to Stop 1 is a gravel road at the west end of a guardrail along the south side of Highway 375, 2.2 mi (3.5 km) by road west of Hancock Summit. The gravel road descends into an arroyo and highway maintenance gravel pit that bounds the northwest side of the mountain. The road continues southwest down the arroyo. The Stop 1 traverse begins at the lowest beds exposed along the bank, 0.5 mi (0.8 km) from Highway 375 (Fig. 12). Coordinates at the Stop 1 parking area and base of traverse: lat 37°24′36.04″N., long 115°24′05.42″W., Crescent Reservoir 7.5′ quadrangle. Rock Units Exposed Simonson Dolostone (brown cliff member, upper alternating member, informal “upper coarse crystalline member”), Fox Mountain Formation, Guilmette Formation, and possibly West Range Limestone (Figs. 5 and 11). The Sevy Dolostone, cherty argillaceous unit, Oxyoke Canyon Sandstone, and lower members of the Simonson Dolostone (Fig. 11) are not exposed along the traverse of Stop 1; see Stop 2. The Simonson was divided into four informal members by Osmond (1954; Figures 5, 11): coarse member, now termed the coarse crystalline member (Johnson et al., 1989), which appears to be transitional from the Oxyoke Canyon Sandstone below and becomes more coarsely crystalline upward; lower alternating member, which contains alternating light and dark bands of beds and grades upward into the brown cliff member, which in turn grades into the upper alternating member. The upper member also commonly becomes coarsely crystalline upward, exhibiting an interval that is informally referred to as the “upper coarse crystalline member.” The Simonson Dolostone exhibits upward-shallowing cycles that are partially exposed in each member. The more dolomitic Simonson cycles are generally thinner and represent shallower conditions than those that are so well preserved in the overlying Guilmette Formation, and thus were probably situated more landward. However, structureless intervals at the base of many cycles probably represent lower intertidal to subtidal bioturbated limestone that was later dolomitized along with the entire formation. The underlying Sevy Dolostone, exposed in ranges nearby, may represent even shallower conditions, probably accumulated more dolomite-prone beds, and is even more strongly altered by regional dolomitization. The Fox Mountain interval was regarded as the upper unit of the Simonson Dolostone in Nevada or the lower member of the

Guilmette Formation in Utah until it was recognized as a regionally widespread unit and given formation status by Sandberg et al. (1997), who divided it into lower and upper members. Part of the upper member at this locality is shown in the outcrop photograph of Figure 13 and the stratigraphic column of Figure 14. The history of studies and partitioning of the Guilmette Formation has already been recounted under “Devonian Formations.” The lower part of the Guilmette on the Stop 1 traverse was divided by Sandberg et al. (1997) into the four intervals shown on Figure 14: the “yellow slope-forming” and “carbonate-platform facies,” the type Alamo Breccia Member, and the “slope facies.” The “slope facies” is only the lower part of the upper Guilmette, which continues upward for an additional ~440 m and contains thick quartz sandstone intervals as well as beds that evidence both deep and shallow carbonate-platform environments. Objectives • Introduce several of the Devonian formations exposed in the ﬁeld guide area. • Describe shallowing-upward cycles and their character in different formations. • Document fossils and ichnofossils of platform facies. • Discover signiﬁcant surfaces; compare with established formation and member limits. • Note trend upward from dolostone of Simonson to mainly limestone of Fox Mountain and limestone/dolostone cycles in the Guilmette and discuss probable controls. • Discuss sedimentary structures and provenance of quartzose sandstone in upper member of Guilmette. • Discover varied character of Devonian exposure surfaces. • Introduce Alamo Breccia and its bolide impact signatures. Traverse The Stop 1 traverse begins in the brown cliff member of the Simonson Dolostone, at the base of the stratigraphic section exposed along the arroyo, and continues eastward up through the stratigraphic section for ~1 mi (~1.6 km) along a ridge that forms the drainage divide along the crest of “Downdropped Mountain” (Fig. 12). The traverse ends at the east end of “Downdropped Mountain,” at the base of the dip slope that marks the top of the exposed platform facies of the Guilmette Formation, or possibly within, or at the top of, a thin interval of West Range Limestone. East of “Downdropped Mountain” is a strike valley of Pilot Shale, accessible by a rough dirt road that enters from the north, and a ridge of Joana Limestone that is cut off at the south end by a major fault. Beyond the fault is the stratigraphic section studied by Estes-Jackson (1996), shown in Figure 2, and a twin of the section traversed at Stop 1. Observations Simonson Dolostone Brown Cliff Member: Rich brown color; conversion to dolomite of varying crystal sizes; relatively open platform evidenced

Devonian carbonate platform of eastern Nevada by several forms of abundant stromatoporoids, some corals, and other invertebrates; cycle boundaries vague; transition upward to upper alternating member. Upper Alternating Member: Trend upward to light-gray color; cycles marked by structureless (bioturbated) lower intervals and algal-laminated upper intervals; fossils less abundant, mainly bulbous stromatoporoids; trend upward to coarser grained dolostone; pockets of yellow-weathering, coarse-grained dolomite crystals interpreted as karst ﬁllings; abrupt shift across bedding upward to ﬁner grained dolostone and absence of yellow crystals, indicating exposure surface and underlying episode of diagenetic recrystallization and solution; altered intervals with zebra rock and coarse crystallization associated with Devonian exposure surfaces and also with much later (Cenozoic) fault zones. Fox Mountain Formation The Fox Mountain was divided into lower and upper members by Sandberg et al. (1997). The contact between those members represents a regional shift from the underlying shallow-platform dolostone formations to open-marine platform limestones, and signals the termination of the long-lived very shallow platform that had existed from the beginning of Devonian deposition. The karst zone under the base of the lower member separates underlying Simonson dolostones from Fox Mountain beds interpreted as peritidal, restricted-marine, and evaporite-solution-breccia limestones. A second karst at the top of lower member separates these beds from open-marine crinoidal wackestones and encrinites of the upper member, which contains nautiloids, brachiopods (including Stringocephalus), and corals. The upper member represents initiation, in the Middle varcus conodont Zone, of the Taghanic onlap, a marine transgression recognized throughout North America and in parts of Europe (e.g., Johnson et al., 1985; Johnson and Sandberg, 1989). See Figure 14 for thicknesses and details of the upper member. Guilmette Formation Yellow Slope-forming Member: The yellow-weathering interval of ﬁne-grained, dark-gray, silty dolostone beds forms the topographic saddle shown in Figure 13 and is underlain by several meters of stromatolitic dolostone that regionally represent the basal beds of the interval (Fig. 14). Note the columnar and branching stromatolites and yellow-weathering, silty dolostone and dolomitic siltstone beds that are replaced upward by darkgray limestones of the overlying member. Carbonate Platform Facies (Member): This member exhibits excellent examples of shallowing-upward cycles. They show basal transgressive erosion surfaces and sediment lags overlain by a lower interval of dark-gray bioturbated limestone that usually contains bulbous and other forms of stromatoporoids, and may contain other invertebrate fossils and oncolites. The upper part of cycles that are completely preserved becomes increasingly lighter gray, dolomitic, and algal laminated. Alamo Breccia Member: As shown on Figure 14, the base of this Member is marked by the subtle Unit D monomict

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detachment breccia, overlain by Unit C megaclasts composed mainly of previously deposited platform cycles similar to those in the underlying member. Unit B is a chaotic heterolithic breccia with impact lapillistone clasts, overlain by Unit A graded beds with shocked and hematite-studded quartz grains (Warme and Sandberg, 1995, 1996; Morrow et al., 2005) and both local and exotic deformed clasts. The traverse along the drainage divide intercepts the Breccia where Unit B reaches the base of the Breccia and separates two Unit C clasts that are hundreds of meters in length. Unit B contains clasts, tens of meters in length, ﬂoating in a chaotic matrix. The stacked graded beds of Unit A are well exposed on the sloping ledge at the top of the Breccia. Slope Facies (Member): The member over the Alamo Breccia is varied. Beds are commonly mottled, fossiliferous limestones of deeper-water aspect (Warme and Sandberg, 1995, 1996; Sandberg et al., 1997), and exhibit one or more hardgrounds. The interval is partly covered, and ends at the thick quartz sandstone beds that continue upward into the lower part of the upper Guilmette (Fig. 12). Upper Guilmette (Member): The lowest interval of Reso’s upper Guilmette is the 35-m-thick “slope facies” member described above, which is overlain by quartzose sandstones that are broadly channeled, cross bedded, and exhibit biogenic sedimentary structures. The middle part contains biostromes and small bioherms of stromatoporoids, and beds rich in corals, gastropods, and bivalves including megalodonts. Limestones of the upper part are thin bedded and quartzose. They contain abundant biogenic sedimentary structures and at the top may include a thin interval of West Range Limestone. Stop 2. Sixmile Flat: Upper Sevy Dolostone, Cherty Argillaceous Unit(?), Oxyoke Canyon Sandstone, and Coarse Crystalline Member and Lower Alternating Members of Simonson Dolostone Location Stop 2 is situated north of Sixmile Flat, a broad valley east of the Hiko Hills (Figs. 9 and 10). From the Crystal Springs rest area starting point, travel 0.8 mi (1.3 km) east to Highway 93, then east on Highway 93 for 10 mi (16 km), past the south end of the Hiko Hills, entering Sixmile Flat and continuing to a dirt road on the left (north) that is the entrance to a cattle pen. Travel 0.4 mi (0.6 km) WNW to the pen, then north along a straight sectionline fence and gravel road for 5.8 mi (9.3 km) to a fence gate and cattle trail on the left (west) (Fig. 15). Stop 2 traverse begins 0.3 mi (0.5 km) west of the gate and can be approached part way by vehicles with high ground clearance. Coordinates at Stop 2 parking area at fenceline: lat 37°40′28.61″N., long 115°04′59.22″W., Hiko NE 7.5′ quadrangle. Rock Units Exposed Sevy Dolostone, cherty argillaceous unit(?), Oxyoke Canyon Sandstone, and coarse crystalline and lower alternating members of the Simonson Dolostone. The traverse can be con-

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tinued upslope and into the brown cliff member, but is truncated northward by Cenozoic layered volcanic rocks (Fig. 15). Objectives • Overview of exposed formations. • Transition from Sevy Dolostone to local expression of cherty argillaceous unit (Dox1? of Fig. 11) and Oxyoke Canyon Sandstone (Dox2 of Fig. 11). • Deeper facies of Oxyoke Canyon Sandstone (Dox2 of Fig. 11): hummocky cross bedding and trace fossils. • Transition to cyclic bedding and gradual coarsening to top of coarse crystalline member. • Karst signatures at top of coarse crystalline member and ﬁner-crystalline dolostone at base of lower alternating member. Traverse From dirt road proceed west on cattle trail to highest point of ﬁrst shallow saddle, then walk south to the upper part of the Sevy Dolostone and start of northward traverse. Observations Cryptic evidence for platform cycles in heavily dolomitized Sevy Dolostone; disseminated ﬁne-grained quartz in upper Sevy (Dox1?) sequence of Figure 11; varied quartz concentrations in upper Oxyoke Canyon (Dox2) sequence of Figure 11; identiﬁcation of ichnotaxa in Oxyoke Canyon; conﬁrmation of hummocky cross-stratiﬁcation (HCS); nature of cycles, increasing coarse crystallinity, and increasing karst features in coarse crystalline member (Dsicxln of Fig. 11); abrupt ﬁning of crystallinity at base of lower alternating member (Dsilalt of Fig. 11). Stop 3. Silver Canyon: Alamo Breccia and “Reso’s Reef” Location From Crystal Springs rest area starting point, travel north 2.5 mi (4.0 km) on Highway 318, into the settlement of Hiko, to entrance of gravel road on left (west; unmarked Logan Canyon Road). Travel northwest for 7 mi (11 km) to “Y” junction; turn right (unmarked Silver Canyon Road). Note excellent views of the white carbonate buildup at top of hill, front right; Alamo Breccia is steep cliff below the buildup. Continue north 1.3 mi (2.1 km) to two small, leveled parking and turnaround areas on right. Coordinates at Stop 3 parking area and base of traverse: lat 37°37′14.10″N., long 115°21′37.74″W., Mount Irish SE 7.5′ quadrangle. Rock Units Exposed Formations exposed north and east of the road are Devonian. To west of the road, in valley, are dark orange-brown-weathering acidic volcanic ﬂows, some marked with American Native petroglyphs. Devonian formations in various fault blocks to east of road can be identiﬁed as Sevy Dolostone below the distinctive twin intervals of dark-brown-weathering Oxyoke Canyon Sandstone,

and coarse crystalline member of Simonson Dolostone above them. Thick brown beds well above the Oxyoke Canyon are in the brown cliff member. Due east, a low ridge in the foreground has almost continuous exposures of upper Simonson Dolostone and Fox Mountain Formation, which are exposed below the distinctive yellow slope-forming member of the Guilmette. The member weathers to a light-yellow band across the near mountain face, and is labeled on Figure 16. The yellow band is overlain by an interval similar to the “carbonate platform facies” of Stop 1 (Fig. 14), labeled as lower Guilmette (Dgl) on Figure 16. This interval is overlain by the cliff-forming Alamo Breccia Member, which extends to the skyline on Figure 16. See description of lower three members of the Guilmette under Stop 1 (Fig. 14). The white carbonate buildup above the Breccia is not visible from the parking location. Objectives • Study graded beds of Unit A, upper interval of Alamo Breccia, containing large clasts of platform carbonates and clasts of impact lapillistone. • Note facies directly over Alamo Breccia that shift west to east, from dark, bedded, off-buildup intervals to massive white carbonate buildup. • Note interﬁngering buildup and off-buildup tongues. • Document evolution of buildup from base to top. • Discuss whether the buildup is mudmound, microbial mound, reef, or other alternatives. • Study karst developed across highest pinnacles of buildup. • Document cyclic shallow-water platform beds over buildup. Traverse Stop 3 traverse is ~1.75 mi (~2.8 km) each way. It begins at the right side of parking areas, northward up a gentle ridge (Fig. 16) that arcs eastward, crests, then drops down to a bench that separates the graded bed Unit A at the top of the Alamo Breccia from the base of the overlying buildup and off-buildup facies. As indicated on Figure 16, the traverse continues eastward around reentrants along the top of the Alamo cliff, to the center of the buildup facies over the Breccia. The last reentrant terminates in a steep wall at the base of the buildup, and exposes the early phases of its development. Return westward to notches where the buildup can be traversed upward to the crest. The traverse ends in dark carbonate beds over the structure. Return by same route, or take shortcut westward to intercept the ridge that leads to the parking area. Observations Beds at the beginning of the traverse dip westward into the Silver Canyon fault zone and are in the upper Guilmette over the Alamo Breccia. They contain low-spired gastropods and large solitary rugose corals. Alamo Breccia The Breccia is exposed along the south-facing slope of the traverse ridge, and the traverse follows the graded Unit A.

Devonian carbonate platform of eastern Nevada

Figure 15. Oblique Google Earth aerial view (eye altitude: 1.8 km) to the west of Stop 2 at Sixmile Flat, showing access road and trail, traverse route, and stratigraphic units including Sevy Dolostone (Dse), lower and upper members of Oxyoke Canyon Sandstone (Dox1? and Dox2, respectively), and coarse crystalline (Dsicxln), lower alternating (Dsilalt), and brown cliff (Dsibc) members of Simonson Dolostone. Image modiﬁed from http://earth.google.com/.

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cess trail Ac

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Mt. Irish

Dgab Dgl Traverse

Dgysf

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Figure 16. Outcrop photograph of southfacing Devonian units east of Silver Canyon (Stop 3), including Simonson Dolostone (Dsi), Fox Mountain Formation (Dfm), and Guilmette Formation (Dgysf, yellow slope-forming member; Dgl, lower member; Dgab, Alamo Breccia Member; and Dgu, upper member, with “Reso’s Reef” organic buildup directly overlying Alamo Breccia). Faulted blocks of Sevy Dolostone (Dse) capped by dark-orange weathered cliffs of Oxyoke Canyon Sandstone (Dox) are visible in foreground. Stop 3 traverse route is shown by dashed white line. The traverse ascends ridgeline from the parking area, out of view to left (west), then follows the top of Alamo Breccia cliff and base of the buildup to the center of the buildup, then ascends the buildup onto the platform cyclic beds over it.

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Scattered clasts of lapillistone may be encountered along the Unit A traverse to the top of the hill. Below the crest, the west-facing Breccia cliff exhibits a large-scale folded clast, tens of meters long. The south-facing Breccia cliff exhibits Unit A graded beds with sparse lapillistone clasts, and is interrupted by several bedparallel or tilted clasts, 10 m or more in length. One ~30-m-long tilted clast protrudes from the graded bed at the top of Unit A. Eastward, toward the carbonate mound, beds over the top of the Breccia become light gray to white, fossiliferous, and packed with stromatoporoids of both hemispherical and elongate forms; some elongate tabular forms are 1 m or more in maximum length. The sheer cliff at the end of the traverse exhibits the graded beds in Unit A at the top of the Breccia, is near the center of the buildup, and displays abundant stromatoporoids and an erect branching coral ~70 cm high, preserved in life position. Carbonate Buildups We use the non-genetic term “buildup” for carbonate accumulations with positive relief, because they exhibit a spectrum of compositions and frameworks and have been differently classiﬁed and variously named in numerous schemes. They include “reefs” that some workers restrict to shallow and wave-resistant bio-accumulations, and deep-water structures that are commonly termed “mudmounds.” A review of reef versus mudmound concepts is beyond the scope of this ﬁeld guide. Discussion papers continue to appear. The volume edited by Monty et al. (1995) brings forth the controversies, and includes the contribution by Pratt (1995) who outlined the wide variability of “mudmounds” in time and space. They occur as early as Proterozoic and in every Phanerozoic period, contain variable proportions of mud-sized particles and potential frameworks, and encompass a variety of recognizable invertebrates that change over geologic time and that may play active or passive roles in mound formation. Most structures contain ﬁne-grained cement that, without detailed petrographic study, is commonly difﬁcult to differentiate from originally ﬁne-grained clastic calcareous particles. Reso’s Reef The buildup at Stop 3 three has been termed a mudmound (e.g., Sandberg et al., 1997), but Dunn (1979) documented potential framework of several species of stromatoporoids and corals in the core of the structure, and used the term of bioherm. She listed invertebrates that include gastropods, brachiopods, and crinoid columnals within and adjacent to the buildup, which appears to contain a more abundant fauna than has been observed beyond its margins. One nautiloid has been found. These components are difﬁcult to observe in the central core facies because of pervasive recrystallization. “Reso’s reef” could beneﬁt from further investigation. An upward traverse through the buildup offers the chance to discuss its genesis: reef versus mudmound or other options. The upper ~10 m of the buildup shows increasing vugs and seams that are stained yellow, red, and orange, indicating that the top was exposed in Devonian time. A karst cave, 1 m in

diameter, near the top exhibits a ﬂoor of bedded sediment that dips at about the same angle as the regional dip, providing a geopetal indicator and evidence that the solution occurred before regional tilting and not after current exposure. Beds over the buildup are cyclic and contain abundant Amphipora, indicating continued platform conditions. The stratigraphic column of Figure 11 was measured nearby. The sequence lacks thick sandstone beds, in contrast to the abundant sandstones in the upper Guilmette at Stops 1 and 4. Stop 4. Hiko Hills: Alamo Breccia and Evacuation Structure Location Stop 4 is located at the south end of the Hiko Hills. It is accessed by traveling east 0.8 mi (1.3 km) from the Crystal Springs rest area starting point to the end of Highway 318, turning north, then east on Highway 93 for 1 mi (1.6 km), and then north (left) onto a dirt road. The road skirts the west side of a gravel pit, becomes rough, and heads east-northeastward to the mouth of a small canyon at the base of the range. The canyon leads upward to outcrops of the Alamo Breccia intercalated with the Guilmette Formation. Coordinates at the Stop 4 parking area and base of traverse: lat 37°32′51.27″N., long 115°12′08.38″W., Hiko 7.5′ quadrangle. Rock Units Exposed Guilmette Formation: Alamo Breccia Member; cyclic platform carbonates in lower Guilmette Formation below the Breccia; cyclic carbonates, carbonate buildups, and broadly channeled sandstones above the Breccia. Objectives • Characterize internal fabric and components of Alamo Breccia. • Note character and distribution of lapillistone within Breccia. • Observe post–Alamo Event recolonization of carbonate platform: fossils, trace fossils, carbonate buildups. • View evacuation of matrix under ﬁnger of Unit C megaclast. Traverse From the mouth of the canyon, proceed up the main gully. The closest Breccia section is on a down-faulted block, detached from the mountain front. It contains excellent exposures of the clasts and matrix of Unit B in the middle part and stacked graded beds of Unit A in the upper part. The Guilmette above the Breccia exhibits platform cycles, which are heavily dolomitized and altered near the fault, which terminates the lower traverse. Upward, across the fault, is a more complete second exposure of the Breccia. The base of the Breccia is exposed across the fault. The Breccia may be traversed vertically to overlying beds, or laterally northward for several hundred meters along the west face of the Hiko Hills to observe variations within the Breccia and the overlying and underlying beds.

Devonian carbonate platform of eastern Nevada Observations A climb through the lower Breccia outcrop reveals the chaotic character of Unit B: heterolithic clasts and variable matrix. The interval of Unit A contains three or more graded beds that become ﬁner grained and thinner upward, and shows load, ﬂame, and dewatering structures at bed boundaries. The top of the Breccia merges with a silty interval that contains faint crossbedding and abundant ichnofossils resembling Teichichnus. This section was discussed by Sandberg et al. (1997). Recovery of the platform biota after the Alamo Event is under study (Tapanila and Anderson, 2007). Across the fault, the upper traverse begins with a thick section of chaotic Unit B matrix and heterolithic clasts. As shown in Figure 17, Unit B terminates upward at a surviving ﬁnger of cyclic beds that project eastward from the top of the 80-mthick Unit C clast visible to the west. Pinto and Warme (2008) described the unusual relationships between the very thick Unit C clast, nearly in its original position, and the laterally adjoining matrix of the Alamo Breccia. The ﬁnger of platform beds extends from the upper part of the thick Unit C clast, over ~60 m of Unit B breccia. During the Alamo Event, this ﬁnger initially remained intact, then fractured and parted in two places, allowing well-sorted clasts from the already accumulated overlying Unit A to cascade down the new slots. The structure is interpreted to have formed when some of the Unit B breccia was evacuated from under the ﬁnger. This process is proposed to be part of the ring-forming adjustments of the transient crater in later phases of the Alamo Event (Pinto and Warme, 2008). Stop 5 provides a second example of this process. Northward, the thick Unit C interval of intact cyclic beds can be traversed ~0.25 mi (~400 m) to a vertical fault, beyond which most of the Unit C clast was destroyed and replaced by the chaotic Unit B, which displays a variety of spectacular clasts tens of meters long. The upper 10 m of the Breccia contains a train of numerous lapillistone clasts that appear to be remnants of the same lapillistone bed. Above the Alamo Breccia are discontinuous light-gray-weathering carbonate buildups with abundant stromatoporoids and fewer colonial corals, and stacks of broadly channelized quartzose sandstones, interbedded with cyclic carbonates. Details of the Breccia in this area were documented by Kuehner (1997).

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canyons between the ridges (Fig. 18), where the Alamo Breccia is displayed, are ~0.6 mi (~1 km) from the highway, and can be reached on foot. Coordinates at the Stop 5 parking area at Highway 93 and base of traverse (approximate): lat 37°04′35.01″N., long 114°59′25.24″W., Pahranagat Wash 7.5′ quadrangle. Rock Units Exposed The lower Guilmette Formation dips more steeply than the structural dip along the western face of the southern end of the Delamar Mountains, so that the Alamo Breccia caps several isolated east-west–trending ridges that slope westward, the underlying carbonate-platform facies member is exposed in the intervening small valleys, and the yellow slope-forming member (terminology of Sandberg et al., 1997) forms an irregular strike valley east of the ridges. Northward, Middle Devonian beds under the yellow slope-forming member crop out along a series of very irregular ridges and valleys. Objectives • Note relatively thin interval of lower Guilmette members. • Observe the scale and distribution of giant tabular clasts that compose ~80% of the Alamo Breccia. Discuss origin and signiﬁcance for genesis of Alamo Breccia. • Note unusually well-sorted matrix between clasts, and lithologic similarity with Unit A at top of Breccia. • Inspect small stromatoporoid patch reef within Unit C detached clast. • Note subtle expression of Unit D monomict detachment breccia. • Time permitting, traverse eastward, down section, to the yellow slope-forming member, and continue northeast into the underlying Middle Devonian formation with abundant bioherms and biostromes of pentamerid brachiopods, stromatoporoids, and corals. Traverse From Highway 93, approach the front of the Delamar Mountains (Fig. 18) and traverse one or more of the isolated ridges of Alamo Breccia and one or more intervening valleys. Time permitting, walk northward in the yellow-weathering strike valley east of the ridges, then northeastward topographically up and stratigraphically down into the exposed Middle Devonian fossilrich dolostone beds.

Stop 5. Southern Delamar Mountains: Stacked Tabular Clasts in Alamo Breccia

Observations

Location Stop 5 is located near the southern end of the Delamar Mountains, where they verge closest to Highway 93 (Fig. 18). From the starting point at the Crystal Springs rest area, travel east 0.8 mi (1.3 km) to the end of Highway 375, then south on Highway 93 for 35 mi (56 km), passing through the settlements of Ash Springs and Alamo. The Alamo Breccia is the irregular interval that caps several ridges that dip westward toward the highway. Entrances to

The lower Guilmette under the Alamo Breccia is similar, but overall thinner, to that at Stops 1, 3, 4, and other locations in the ﬁeld trip area to the north. However, the fossil-rich Middle Devonian interval, under the yellow slope-forming member, may be either the brown cliff member of the Simonson Dolostone or the upper member of the Fox Mountain Formation. This stratigraphy suggests that the upper alternating member of the Simonson and the lower member of the Fox Mountain were

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Figure 17. Photograph and diagram of Alamo Breccia evacuation structure at the south end of the Hiko Hills (Stop 4). Breccia interval is 100 m thick. (A) View to northeast showing characteristic shallow-water carbonate-platform cyclic bedding of the Guilmette Formation above and below Breccia. (B) Left: unusually thick (~80 m) preserved Unit C megaclast with thin (~10 m) Unit A graded beds over top. Center and right: cyclic carbonates equivalent to the megaclast disintegrated down to the level of detachment (Unit D), except for an upper ﬁnger of beds (F) that extended over newly formed Unit B chaotic breccia. The ﬁnger broke into pieces as some of the underlying Unit B was evacuated. Lapillistone clasts (L) as much as 50 m under the ﬁnger represent early-precipitated beds that were broken and mixed with Unit B. One or more later lapillistone beds precipitated from the impact plume, became partially lithiﬁed, and were preserved as a discontinuous trail of broken and smeared lapillistone within a graded bed of Unit A (black circles), which formed across the whole area in a late phase of the Alamo Event after the ﬁnger collapsed. Post-Event deposits in this area include buildups, up to ~40 m high, containing stromatoporoids and corals (R). From Pinto and Warme (2008).

Devonian carbonate platform of eastern Nevada

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Alamo Breccia outcrops Tv

Gr e at Ba sin Hw y

Figure 18. Oblique Google Earth aerial view (eye altitude: 1.2 km) to the southeast of Stop 5 at the southern Delamar Mountains, showing isolated Alamo Breccia outcrops (labeled) capping eastwest–trending ridges. Stacked tabular clasts within the Breccia are visible. Highway 93 is labeled. Tv indicates Tertiary volcanic rocks. Image modiﬁed from http://earth.google.com/.

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either not deposited or were eroded prior to deposition of the yellow slope-forming member. The Alamo Breccia at Stop 5 is atypical and difﬁcult to interpret as part of the Alamo Breccia scenario. The long, tabular, stacked clasts and sparse, presorted interclast matrix (Fig. 19) have not been observed at other Alamo Breccia localities. Stop 5 is ~85 km from the present position of Tempiute Mountain (unrestored), Stop 6, where the beds are interpreted as a fragment of the Alamo crater rim. At Stop 5, the Breccia is ~60 m thick and contains stacked irregular clasts that are a hundred meters or more in length. The mega-fabric of the clasts suggests that they moved eastward, not northwestward in the direction of Tempiute Mountain (Figs. 9 and 20) or westward toward the paleo-platform margin. The orientation of large-scale clasts in the Alamo Breccia at other localities, much closer to Tempiute Mountain, indicates movement approximately westward, as if they were sliding toward the crater or the seaward platform margin (Kuehner, 1997; Warme and Kuehner, 1998). In addition, the total thickness of the Alamo Breccia here is ~60 m (Fig. 19), whereas it is thinner at some localities closer to the present position of Tempiute Mountain or to the trend of the platform margin. Pinto and Warme (2008) proposed that the clasts at Stop 5 were detached, moved eastward, and stacked over one another during the ring-forming processes associated with adjustment of the transient crater. In this scenario, listric-fault-bounded, arcuate segments of the platform dipped outward, fault footwalls

became rings, and near surface intervals of beds detached and slid away, perhaps collecting near the footwall of the next ring fault outward. Figure 9 shows that both Stops 4 and 5 are located near the outer edge of the Ring Realm of the Alamo Breccia, suggesting that signiﬁcant faulting formed one or more outer crater rings and an adjacent moat that collected Alamo Breccia. Stop 6. Tempiute Mountain: Crater-Rim Impact Stratigraphy Location From the Crystal Springs rest area travel ~33 mi (53 km) on Highway 375 westward, past Stop 1, then northwest to the Coyote Summit road marker. Continue northwest on Highway 375 for 3.0 mi (4.8 km) and turn east on a dirt road that leads to Tempiute Mountain (Fig. 20). Travel eastward for 1.4 mi (2.2 km) to fork; turn left (northeast) and continue 1.7 mi (2.7 km) to canyon mouth, where road is washed out. Continue on foot eastward up canyon through old mining works to cliff of Alamo Breccia, Stop 6A (Fig. 20). Coordinates at the Stop 6 parking area and base of traverse: lat 37°36′58.44″N., long 115°38′51.84″W., Tempiute Mountain South 7.5′ quadrangle. Rock Units Exposed The west face of Tempiute Mountain clearly exposes Paleozoic formations from the Ordovician Eureka Quartzite to the

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Figure 19. View to south-southwest of Alamo Breccia outcrop that caps the dip slope at the southern end of the Delamar Mountains (Stop 5). Image is tilted to restore horizontal, which is indicated by trend of ridge of the Sheep Range in distance at right in (A) and slanted line in (B). (A) Photo of northern slope of one of several east-west–trending ridges that display Alamo Breccia with unusual tabular megaclasts and very little matrix. Note 60-m scale of total Breccia. (B) Units of the Breccia include the detachment surface (D), jumbled megaclasts that continue to the top of the Breccia in most places (C), limited internal chaotic Breccia (B), and topmost graded beds that are now in the process of eroding away (A). R represents stromatoporoid buildups within displaced clasts. From Pinto and Warme (2008).

Mj Dp

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Figure 20. Oblique Google Earth aerial view (eye altitude: 2.1 km) to the southeast of Stop 6 at Tempiute Mountain, showing access road, traverse route, Stop 6A and 6B locations, and stratigraphic units including Eureka Quartzite (Oe), Ely Springs Dolostone (Oes), Laketown Dolostone (Sl), Sevy Dolostone (Dse), Oxyoke Canyon Sandstone (Dox), Sentinel Mountain Dolostone (Dsm), Alamo Breccia (Dab), Devils Gate Limestone (Ddg), Pilot Shale (MDp), and Joana Limestone (Mj). The detailed, complex, crater-rim stratigraphy of the Alamo Breccia and adjacent units is depicted in the columnar section of Figure 21. Grants Peak (7199 ft, 2182 m) and Wildcat Canyon are also indicated. Image modiﬁed from http:// earth.google.com/.

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Devonian carbonate platform of eastern Nevada Mississippian Joana Limestone (Fig. 20). Devonian formations include the Sevy Dolostone, cherty argillaceous unit, Oxyoke Canyon Sandstone, Sentinel Mountain Dolostone (and Bay State Dolostone, if it was locally preserved), Alamo Breccia, Devils Gate Limestone, and Devonian to Mississippian Pilot Shale. The Sentinel Mountain and Bay State Dolostones are deeper platform equivalents of the Simonson Dolostone. All of the Guilmette Formation, or its equivalent that was deposited in this section, and probably all of the Fox Mountain Formation, were cut out and replaced by newly deﬁned Units 3–5 of the Alamo Breccia, as shown in Figure 21 and so far discovered only at Tempiute Mountain. Objectives • Review characteristics of impact beds, interpreted as preserved inner rim of Alamo crater, as depicted in the stratigraphic column of Figure 21. • Stop 6A: Note features of Unit 3, interpreted as large-scale fallback breccia clasts and matrix, and base of resurge beds, Unit 5. • Stop 6B: Discriminate between characteristics of parautochthonous impact breccias (Unit 1), developed nearly in situ in Lower to Middle Devonian Oxyoke Canyon Sandstone and Middle Devonian Sentinel Mountain Dolostone (and Bay State dolostone?) and contrasting exotic clasts characteristic of Units 3–5. • Note features of injected sedimentary sills and dikes, Unit 2, into Unit 1. • Discuss interpretations for shatter-cone-like structures in Unit 1 dolostones. • Note features of Unit 3, interpreted as intensely deformed fallback breccia; Unit 4, interpreted as smeared or ﬂattened fallback breccia; and Unit 5, interpreted as resurge of impact debris into crater, against inner crater rim, or across newly formed slope. • Discuss depositional environment of post-impact deepwater limestones labeled Devils Gate (Fig. 21). Traverse Steeply dipping Paleozoic formations are exposed along the washed-out track in the axis of the canyon, and are labeled on Figure 20. Most noticeable is the Ordovician Eureka Quartzite near the canyon mouth, and the Devonian Oxyoke Canyon Sandstone midway to the cliff of Alamo Breccia that terminates the eastward segment of the road, which is Stop 6A (Fig. 20). Stop 6B requires a signiﬁcant commitment of time. Return west ~330 ft (~100 m) to intersection of steep road that leads northward, up past mine workings and an abandoned shaft, to drainage divide and a view northward of the stratigraphic section on Tempiute Mountain, northwestward across Sand Spring Valley, and west to the settlement of Rachel on Highway 375. This point is the beginning of the traverse for Stop 6B, which runs along the ridge that leads eastward to the range crest (Fig. 20). Return via the same route.

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Observations Stop 6A Exposures along the north side of the easternmost ~165 ft (~50 m) of road show 20 or more lithologies of smashed and interpenetrated Unit 3 megaclasts, some as much as several meters across, that are interpreted as a heterolithic fallback Alamo Breccia (Pinto and Warme, 2008). Near the cliff face at the end of the road is a sandstone clast identiﬁed as the only megascopic fragment of probable Ordovician Eureka Quartzite found thus far, providing evidence for crater excavation to depths of 1.5–2.0 km. Other clasts appear milky white and marbled, and are candidates for partial melting of target rock during the cratering process (Pinto, 2006). The fallback breccia matrix is best exposed in sparse outcrops on the steep slope that leads to the next ridge north. The matrix contains a variety of clasts, shocked quartz grains, and impact lapilli mixed with the heterolithic clasts (Pinto and Warme, 2008), Stop 6B From the Stop B starting point (Fig. 20) the road becomes level and continues a short distance northward, then eastward into the next east-west drainage, and provides excellent views of the next ridge to the north and its south-facing foreground slope. This excursion of the traverse is shown on Figure 20. Visible from west to east, in ascending order, are: the irregularly siliciﬁed upper Sevy Dolostone; the well-bedded cherty argillaceous unit that is not well represented in localities to the east but is a deeper-water facies that may correlate to Dox1 of Figure 11; indurated and cross-bedded Oxyoke Canyon Sandstone, equivalent to Dox2 of Figure 11; and Sentinel Mountain Dolostone and possibly Bay State Dolostone, equivalent to the Simonson Dolostone of shallower facies on the platform. Weathered brownish-orange patches in the Sentinel Mountain are injected and sediment-ﬁlled sills and dikes of Unit 2, which can be inspected in outcrops near the canyon axis at the end of the road (Fig. 20). The east end of the ridge contains the Alamo Breccia, Devils Gate Limestone, Pilot Shale, and Joana Limestone, which cannot be discriminated from this viewpoint, but can be traversed along the ridge of Stop 6B (Fig. 20). Return to Stop 6B starting point. Beginning at the Stop 6B starting point (Fig. 20), the traverse eastward along the ridge axis exhibits almost continuous exposures that represent impact stratigraphy, as interpreted by Pinto (2006) and Pinto and Warme (2008). The traverse crosses breccias of parautochthonous deformed bedrock overlain by exotic fallback and resurge clasts, all interpreted to be preserved on the inner rim of the transient Alamo crater (Fig. 21). Along the ridge, the cross-bedded, silica-cemented, upper part of the Oxyoke Canyon Sandstone (S1A), together with the very dark-gray to black beds of the Sentinel Mountain Dolostone (S1B), are designated as the parautochthonous Unit 1 impact breccias on the stratigraphic column of Figure 21. This 300-m-thick interval is thoroughly fractured, faulted, and injected with the sediment-ﬁlled

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Figure 21. Columnar section of Devonian rocks at Tempiute Mountain (Stop 6), showing the thickness and character of Alamo impactogenic Units 1–5, the underlying upper part of the Sevy Dolostone (Dse) and cherty argillaceous unit, the overlying deep-water Devils Gate Limestone (Ddg), and the Pilot Shale (MDp). Parautochthonous impact breccias of Unit 1 are composed of the Oxyoke Canyon Sandstone (S1A) and Sentinel Mountain Dolostone (S1B). Unit 2 is injected polymict breccia that tends to be thickest where it swells, pinches and separates S1A and S1B. S1B contains shatter-cone–like structures, interpreted as impact-induced. Units 3 and 4 are composed of allogenic limestone clasts with varied matrices that bear shocked quartz, lapilli, and other particles interpreted as shock and melt indicators. Unit 5 is two thick resurge breccias. From Pinto and Warme (2008).

Devonian carbonate platform of eastern Nevada dike-and-sill system of Unit 2 and interpreted as nearly in situ altered bedrock. Note the abundant shatter-cone–like structures and related features in the black dolostones. The dolostone breccias of Unit 1 are overlain by the mainly limestone breccia Units 3 and 4. Unit 3 is a well-exposed, roughly graded, fallback breccia with clasts as much 10 m in exposed length. The better sorted smaller clasts of Unit 4, averaging ~5 cm in length, are interpreted as a separate fallback event of heterolithic clasts that were partially melted, weakened, and smeared. In contrast, Unit 5 is composed of two 30-m-thick graded beds of generally angular, clean-washed, heterolithic clasts that are interpreted as resurge ﬂows deposited against the inner wall of the crater rim or across a new slope that formed by collapse of the platform margin during later stages of the Alamo Event. The upper graded bed ﬁnes imperceptibly into the post-Breccia deep-water limestones labeled Devils Gate on Figure 21. Of interest in the Devils Gate is the lack of benthic invertebrates and ichnotaxa, presence of small pelagic or nektonic species, and several debrites, as much as ~4 m thick, of locally derived slumps and ripup clasts with varying proportions of quartz grains derived from elsewhere. Hypotheses for the depositional environment of the Devils Gate at this locality include post-Event crater ﬁll and slope veneer. ACKNOWLEDGMENTS We thank Wanda Taylor for beneﬁcial review of the manuscript, and Gene Smith and Ernest Duebendorfer for their editorial handling of the paper. Jesús Pinto provided data and assistance with ﬁgures. Morrow’s contribution to this manuscript is based upon work supported by the National Science Foundation under Grant No. #0518166. Sandberg’s contribution is supported by the U.S. Geological Survey Bradley Scholar Program. REFERENCES CITED Arthur, M.A., and Garrison, R.E., 1986, Milankovitch cycles through geologic time: Paleoceanography, v. 1, p. 355–586. Burgess, P.M., 2006, The signal and the noise: forward modeling of allocyclic and autocyclic processes inﬂuencing peritidal carbonate stacking patterns: Journal of Sedimentary Research, v. 76, p. 962–977. Chamberlain, A.K., 1999, Structure and Devonian stratigraphy of the Timpahute Range, Nevada [Ph.D. thesis]: Golden, Colorado School of Mines, 564 p. Chamberlain, A.K., and Warme, J.E., 1996, Devonian sequences and sequence boundaries, Timpahute Range, Nevada, in Longman, M.W., and Sonnenfeld, M.D., eds., Paleozoic systems of the Rocky Mountain region: Rocky Mountain Section, SEPM (Society for Sedimentary Geology), p. 63–84. Cook, H.E., and Corboy, J.J., 2004, Great Basin Paleozoic carbonate platform: Facies, facies transitions, depositional models, platform architecture, sequence stratigraphy, and predictive mineral host models: U.S. Geological Survey Open-File Report 2004-1078, 129 p. Cook, H.E., Hine, A.C., and Mullins, H.T., editors, 1983, Platform margin and deep water carbonates: Society of Economic Paleontologists and Mineralogists Lecture Notes for Short Course No. 12, 573 p. Crafford, A.E.J., and Grauch, V.J.S., 2002, Geologic and geophysical evidence for the inﬂuence of deep crustal structures on Paleozoic tectonics and the alignment of world-class gold deposits, north-central Nevada, USA: Ore Geology Reviews, v. 21, p. 157–184. Crevello, P.D., Wilson, J.L., Sarg, J.F., and Read, J.F., eds., 1989, Controls on carbonate platform and basin development: Society of Economic Paleontologists and Mineralogists Special Publication 44, 405 p.

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Devonian carbonate platform of eastern Nevada and Fritsche, A.E., eds., Paleozoic paleogeography of the western United States, Paciﬁc Coast Paleogeography Symposium I: Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, p. 67–85. Poole, F.G., and Sandberg, C.A., 1991, Mississippian paleogeography and conodont biostratigraphy of the western United States, in Cooper, J.D., and Stevens, C.H., eds., Paleozoic paleogeography of the western United States II: Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, v. 67, p. 107–136. Poole, F.G., Sandberg, C.A., and Boucot, A.J., 1977, Silurian and Devonian paleogeography and of the western United States, in Stewart, J.H., Stevens, C.H., and Fritsche, A.E., eds., Paleozoic paleogeography of the western United States, Paciﬁc Coast Paleogeography Symposium I: Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, p. 39–65. Poole, F.G., Stewart, J.H., Palmer, A.R., Sandberg, C.A., Madrid, R.J., Ross, R.J., Jr., Hintze, L.F., Miller, M.M., and Wrucke, C.T., 1992, Latest Precambrian to latest Devonian time; Development of a continental margin, in Burchﬁel, B.C., Lipman, P.W., and Zoback, M.L., eds., The Cordilleran Orogen: Conterminous U.S.: Boulder, Colorado, Geological Society of America, The Geology of North America, v. G-3, p. 9–56. Pratt, B.R., 1995, The origin, biota and evolution of deep-water mud-mounds, in Monty, C.L.V., Bosence, D.W.J., Bridges, P.H., and Pratt, B.R., eds., Carbonate mud-mounds, their origin and evolution: International Association of Sedimentologists Special Publication No. 23: Oxford, Blackwell Science, p. 49–123. Read, J.F., 1982, Carbonate platforms of passive (extensional) continental margins: Types, characteristics and evolution: Tectonophysics, v. 81, p. 195– 212, doi: 10.1016/0040-1951(82)90129-9. Reso, A., 1963, Composite columnar section of exposed Paleozoic and Cenozoic rocks in the Pahranagat Range, Lincoln County, Nevada: Geological Society of America Bulletin, v. 74, p. 901–918, doi: 10.1130/00167606(1963)74[901:CCSOEP]2.0.CO;2. Roberts, R.J., Hotz, P.E., Gilluly, J., and Ferguson, H.G., 1958, Paleozoic rocks of north-central Nevada: The American Association of Petroleum Geologists Bulletin, v. 42, no. 12, p. 2813–2857. Ross, R.J., Jr., 1977, Ordovician paleogeography of the western United States, in Stewart, J.H., Stevens, C.H., and Fritsche, A.E., eds., Paleozoic paleogeography of the western United States, Paciﬁc Coast Paleogeography Symposium I: Paciﬁc Section, Society of Economic Paleontologists and Mineralogists, p. 19–38. Sandberg, C.A., and Poole, F.G., 1977, Conodont biostratigraphy and depositional complexes of Upper Devonian cratonic-platform and continentalshelf rocks in the western United States, in Murphy, M.A., Berry, W.B.N., and Sandberg, C.A., eds., Western North America: Devonian, Proceedings of the 1977 Annual Meeting of The Paleontological Society: Riverside, University of California, Campus Museum Contribution 4, p. 144–182. Sandberg, C.A., and Ziegler, W., 1998, Comments on proposed Frasnian and Famennian subdivisions: Document submitted to IUGS Subcommission on Devonian Stratigraphy, Bologna, Italy, June 22, 1998, 6 p. Sandberg, C.A., Ziegler, W., Dreesen, R., and Butler, J.L., 1988, Part 3: Late Frasnian mass extinction: Conodont event stratigraphy, global changes, and possible causes: Courier Forschungsinstitut Senckenberg, v. 102, p. 263–307. Sandberg, C.A., Poole, F.G., and Johnson, J.G., 1989, Upper Devonian of western United States, in McMillan, N.J., Embry, A.F., and Glass, D.J., eds., Devonian of the World: Canadian Society of Petroleum Geologists, Memoir 14, v. I, Regional Syntheses, p. 183–220. Sandberg, C.A., Morrow, J.R., and Warme, J.E., 1997, Late Devonian Alamo Impact Event, global Kellwasser Events, and major eustatic events, eastern Great Basin, Nevada and Utah: Brigham Young University Geology Studies, v. 42, no. pt. 1, p. 129–160. Sandberg, C.A., Morrow, J.R., and Ziegler, W., 2002, Late Devonian sea-level changes, catastrophic events, and mass extinctions, in Koeberl, C., and MacLeod, K.G., eds., Catastrophic events and mass extinctions: Impacts and beyond: Geological Society of America Special Paper 356, p. 473–487. Sandberg, C.A., Morrow, J.R., Poole, F.G., and Ziegler, W., 2003, Middle Devonian to Early Carboniferous event stratigraphy of Devils Gate and Northern Antelope Range sections, Nevada, U.S.A: Courier Forschungsinstitut Senckenberg, v. 242, p. 187–207. Sandberg, C.A., Poole, F.G., and Morrow, J.R., 2005, Milk Spring channels provide further evidence of oceanic, >1.7-km-deep Late Devonian Alamo crater, southern Nevada [abs.]: Lunar and Planetary Institute Contribution No. 1234, abstract 1538, 2 p. Sandberg, C.A., Poole, F.G., and Morrow, J.R., 2006, Reinterpretation of Milk Spring and other oceanic deposits resulting from 382-Ma Alamo impact, southern Nevada, U.S.A., in Papers Presented to Conference on Impact

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The Geological Society of America Field Guide 11 2008

Dinosaurs and dunes! Sedimentology and paleontology of the Mesozoic in the Valley of Fire State Park Joshua W. Bonde* Department of Geoscience, University of Nevada, Las Vegas, 4505 S. Maryland Parkway, Las Vegas, Nevada 89154, USA David J. Varricchio Frankie D. Jackson Department of Earth Sciences, Montana State University, Bozeman, Montana 59717-3480, USA David B. Loope Department of Geosciences, 322 Bessey Hall, University of Nebraska, Lincoln, Nebraska 68588, USA Aubrey M. Shirk Department of Geoscience, University of Nevada, Las Vegas, 4505 S. Maryland Parkway, Las Vegas, Nevada 89154, USA

ABSTRACT This ﬁeld trip covers sedimentological and paleontological research being conducted on the Jurassic Aztec Sandstone and Lower Cretaceous Willow Tank Formation in Valley of Fire State Park. Valley of Fire State Park is located in southern Nevada, just outside of the town of Overton. The Jurassic Aztec Sandstone is equivalent to the Navajo and Nugget Sandstones; these formations together record an aerially large erg complex along the western margin of North America during the time of deposition. Invertebrate and vertebrate ichnofossils are not uncommon in portions of these Jurassic formations. The Willow Tank Formation is composed of the deposits of both a braided and anastomosed ﬂuvial system. This system drained off the paleohigh of the Sevier fold and thrust front to the west, during Early Cretaceous time. Recently a diverse vertebrate assemblage has been discovered from this formation. The fauna of the Willow Tank Formation are similar to other Early Cretaceous faunas from western North America. The vertebrate remains recovered include three taxa of ﬁsh, three to four taxa of turtle, crocodilian, iguanodontian, thyreophoran, dromaeosaur, tyrannosauroid, two theropod ootaxa, and a titanosauriform. In addition to the vertebrate elements, two fern morphotypes have been found. Through the course of this ﬁeld trip participants will see extensive exposures of Aztec Sandstone, including vertebrate ichnofossils. Participants will also hike to vertebrate bearing-beds of the Willow Tank Formation. Keywords: Aztec Sandstone, Willow Tank Formation, Early Cretaceous, Valley of Fire, southern Nevada. *[email protected] Bonde, J.W., Varricchio, D.J., Jackson, F.D., Loope, D.B., and Shirk, A.M., 2008, Dinosaurs and dunes! Sedimentology and paleontology of the Mesozoic in the Valley of Fire State Park, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 249–262, doi: 10.1130/2008.ﬂd011(11). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION This ﬁeld trip and guide provide an overview of the stratigraphic, paleoenvironmental, and paleontologic aspects of Valley of Fire State Park. Located about an hour’s drive northeast of Las Vegas, Valley of Fire is the oldest and largest state park in Nevada (Fig. 1). Due to the sparse vegetation and arid environment of the Mojave Desert, the park provides exceptional views of the Mesozoic stratigraphy of the North Muddy Mountains. Stops on this one-day ﬁeld trip focus on the Aztec Sandstone, a portion of a large erg complex that occupied western North America during the Early Jurassic, and the Willow Tank Formation, which records Early Cretaceous sedimentation in the foredeep of the Sevier fold and thrust front (Schmitt and Kohout, 1986). The deposition of the Willow Tank Formation represents an important time of biotic transition in North America, as endemic faunas yielded to more cosmopolitan ones that incorporated Asian taxa. These latter faunas would characterize the entire Late Cretaceous of North America (Kirkland et al., 1999). Recent discoveries of a diverse vertebrate assemblage in the Willow Tank Formation may provide important information in understanding terrestrial ecosystems during this transitional period in the latest Early Cretaceous. STRATIGRAPHY AND SEDIMENTOLOGY OF THE AZTEC SANDSTONE The Aztec Sandstone exposed in Valley of Fire State Park is Early Jurassic in age, and correlative with the Navajo Sandstone (southern and central Utah, northern Arizona, western Colorado), and the Nugget Sandstone (northernmost Utah and Colorado, Wyoming, eastern Idaho; Kocurek and Dott, 1983). These rocks reach their greatest thickness in southwestern Utah, where the Navajo Sandstone exceeds 700 m (Blakey et al., 1988). Largescale cross-strata prominent in these formations indicate southward and southeastward migration of very large bedforms (Peterson, 1988). Wind ripples superimposed on the ﬂanks of the large dunes generated millimeter-scale laminations; these laminations provide the best evidence for an eolian origin of the sandstones (Hunter, 1977). The Aztec Sandstone and other early Mesozoic rocks in Valley of Fire State Park were probably deposited tens of kilometers to the north and east and moved southwestward during late Tertiary extensional faulting. The similarity of cross-bed dip directions between the Aztec and Navajo Sandstone on the Colorado Plateau indicate little if any rotation of the displaced block about a vertical axis during this movement (Marzolf, 1990). Studies of detrital zircons indicate that westward-ﬂowing rivers provide the primary sources of sand for the giant dune ﬁelds that accumulated near the western margin of Pangaea during Permian and Jurassic time (Dickinson and Gehrels, 2003; Rahl et al., 2003). A large population of 1.1 Ga zircons indicates that much of the sand was derived from the Grenville Belt of the Appalachian Mountains. Most paleomagnetic evidence indicates that southern Utah and northern Arizona lay within 10° of the equator in Early Jurassic time (Steiner, 1983). Climate models show that the

prevailing winds that drove the migrating dunes (as indicated by the dip directions of the sandstones) were trade winds (northern Colorado Plateau) and tropical westerlies (central and southern Plateau; Chandler et al., 1992; Loope et al., 2004). The presence of a supercontinent straddling the equator generated a “megamonsoon” circulation (Kutzbach and Gallimore, 1989); westerlies swept across the paleoequator during December, January, and February when a large low-pressure cell developed in the southern hemisphere. The cross-equatorial winds reversed direction every six months. Southto-north winds (June, July, August) eroded the leeward dune slopes of the Jurassic dunes, generating reactivation surfaces and annual depositional cycles (Hunter and Rubin, 1983). Invertebrate burrows and vertebrate tracks are abundant in portions of the Aztec-Navajo-Nugget Sandstones, suggesting that the paleoclimate was not hyperarid (Loope and Rowe, 2003; Loope, 2006). Siliciﬁed tree trunks and rhizoliths (mineralized plant roots) are present in the Navajo Sandstone of southeastern Utah (Parrish and Falcon-Lang, 2007; Loope, 1988). The positions of slumped cross-strata within Navajo annual depositional cycles near Kanab, Utah, suggest that rain events did not coincide with the time of prevailing winds, and probably fell during JuneJuly-August (Loope et al., 2001). Many Navajo and Aztec exposures display large-scale, softsediment deformation. In a study carried out in the Aztec Sandstone west of Las Vegas, Horowitz (1982) made a strong case that deformation took place during paleoseismic events when the water table lay just below interdune surfaces. The loose packing of avalanche beds (grainﬂow strata of Hunter, 1977) probably made the dune deposits especially vulnerable to liquefaction. Rocks that are dominantly composed of wind-ripple lamination rarely show evidence of liquefaction. Deformation bands (Aydin, 1978) are well-developed in porous sandstones, and record millimeter-scale structural displacements that involve pore-space collapse, grain-scale fracturing, and cataclasis. Deformation bands are prominent in most of the eolian sandstones of western United States (Davis, 1999), and are important barriers to ﬂuid ﬂow in aquifers and hydrocarbon reservoirs. The Aztec Sandstone in the Valley of Fire can be divided into three subunits that correspond to stratigraphic position, outcrop color, and degree of cementation (Flodin et al., 2003). The well-cemented lower subunit appears distinctly red in outcrop. The poorly cemented middle subunit is predominantly buff colored with minor zones of orange, purple, and gray. The orange upper subunit is moderately to poorly cemented (Flodin et al., 2003). Because these color differences do not always correspond with bedding, they likely represent post-deposition, diagenetic changes. Flodin et al. (2003) consider them to have been established prior to initial deformation. SEDIMENTOLOGY AND STRATIGRAPHY OF THE WILLOW TANK FORMATION The mid Cretaceous (Albian) Willow Tank Formation unconformably overlies the Aztec Sandstone and is, in turn, conformably

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Figure 1. (A) Location of Valley of Fire within the state of Nevada. (B) Location of valley of Fire in North Muddy Mountains with ﬁeld trip stops labeled. (C) Location of ﬁeld trip stops within Valley of Fire.

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overlain by the Upper Cretaceous (Cenomanian) Baseline Sandstone (Fig. 2). Exposures of the Willow Tank Formation occur in the Valley of Fire State Park, Bowl of Fire, Gale Hills, and the Virgin Mountains in southern Nevada (Bohannon, 1983). The ﬁrst study of the Willow Tank Formation, conducted by Chester Longwell (1921), combined the Cretaceous and Cenozoic formations into what Longwell termed the Overton Fanglomerate. Longwell (1949) later differentiated the Cenozoic from the Cretaceous formations, further subdividing the Cretaceous section into the Lower Cretaceous Willow Tank Formation and the Upper Cretaceous Baseline Sandstone. In the study area, the Willow Tank Formation strikes ~150° and dips ~35° to the east. The formation is informally divided into two members: a 5–15-m-thick basal conglomerate and a 60–150-m-thick mudrock member (Carpenter, 1989). The Basal Conglomerate Member The basal conglomerate member of the Willow Tank Formation is a prominent ridge-forming stratum. A common element in the Lower Cretaceous formations of the Sevier foreland basin, these conglomerates are thought to represent synorogenic sedimentation associated with either the inception of thrust faulting in the region, or possibly thermal doming prior to thrusting (Heller and Paola, 1989). Clasts included in the basal member of the Willow Tank Formation are predominantly limestone and chert, derived from the Paleozoic rocks in the upper plate of the thrust faults present within the region. The basal member records paleoﬂow from the west (Reese, 1989), thus the exposures likely represent a transect perpendicular to current direction and provide a depositional cross section of the formation. Reese (1989) proposed that the basal conglomerate member of the Willow Tank Formation included two conglomerate bodies, separated by a thin carbonate paleosol. Further, she concluded that the deposits represented deposition in a Donjek-type braidplain (Reese, 1989). Based upon the persistence of limestone clasts in the conglomerate and the carbonate paleosol horizon, Reese (1989) concluded that the basal member was deposited in an arid to semiarid environment. A sharp contact separates the basal conglomerate from the upper member of the Willow Tank Formation. The Upper Mudrock Member The mudrock member consists of thin- to medium-bedded gray claystone and sandstone (Carpenter, 1989). Recent radiometric dates for the upper member produced late Albian ages of 101.6 ± 1 Ma to 99.9 ± 2 Ma (Troyer et al., 2006). Presence of the fern Tempskya also supports a mid Cretaceous age for the Willow Tank Formation (Ash and Read, 1976). Measured sections in the upper member within the study area include ~15% volcanic and 85% siliciclastic lithofacies. Mudrocks represent the most common lithology, accounting for ~75% of the siliciclastic deposits. Most mudrocks are texturally massive, representing overbank ﬁnes; many exhibit mottling, red oxidized color, and contain carbonate nodules (Fig. 3), attributes consistent with paleosol

development (Mack et al., 1993). The abundance of carbonate nodules present in some beds supports an interpretation of these ancient soils as Calcisols (after Mack et al., 1993). The mudrocks are interbedded with two sandstones lithofacies: (1) associated ﬁne- to medium-grained, trough cross-bedded and horizontally bedded sandstones that occur in single story, lenticular beds, and (2) massive to horizontally bedded sandstones with sheet-like geometry. The sheet-like sandstones are far more laterally persistent (100s of meters) than the lenticular sandstones (10s of meters). The lenticular sandstones are interpreted as channel-ﬁll deposits, while the sheet-like sandstones represent crevasse splay sand bodies. Sandstones within the park lack lateral accretion structures and ﬁning upward sequences, typical of deposition in a meandering ﬂuvial system (Allen, 1977; Hallam, 1981). In addition, the lack of coarse-grained (i.e., pebble) clasts and the prevalence of overbank ﬁnes also eliminate deposition in a braided ﬂuvial system (Allen, 1977; Hallam, 1981). Deposits of the upper member of the Willow Tank Formation most likely represent deposition in an anastomosing ﬂuvial system. Previous analyses concluded that the Willow Tank Formation resulted from deposition in the foredeep of the Sevier foreland basin (Schmitt and Kohout, 1986; Schmitt and Aschoff, 2003). A rapidly subsiding basin and a preponderance of overbank muds are often invoked as a cause for the formation of a multi-channeled, vertically accreting ﬂuvial system (Nadon, 1994; Makaske, 2001; Wang et al., 2005a). The single-storied pattern of the channel sandstones becomes more apparent when viewed perpendicular to strike (Fig. 4), a characteristic common to anastomosed ﬂuvial systems (Nadon, 1994; Makaske, 2001). Paleosol attributes (reddening, mottling, carbonate nodules) suggest that deposition occurred in an arid to semiarid environment (Mack et al., 1993; Retallack, 2001). TERRESTRIAL VERTEBRATE FAUNAS OF THE EARLY CRETACEOUS Deposition of the Willow Tank Formation in Valley of Fire State Park occurred during the tectonically active Early Cretaceous period. This 45 million year interval of the Early Cretaceous witnessed the continued separation of Laurasia and Gondwana and the initial break-up of these two supercontinents into smaller landmasses. Sea level rose nearly continuously throughout the Early Cretaceous, eventually reaching a Mesozoic maximum in the early Late Cretaceous (Vail et al., 1977). High sea level, epeiric seas, new circulation routes, and perhaps higher atmospheric CO2 concentrations produced warm, equable climates (Barron, 1983). Within this setting, angiosperms evolved and diversiﬁed and by the latest Early Cretaceous had spread from pole to pole (Feild and Arens, 2005). In addition to ﬂowering plants, all three clades of modern mammals (monotremes, marsupials, and placentals), originated during this same time interval (Archer et al., 1985; Ji et al., 2002; Luo et al., 2003), but consisted primarily of primitive lineages not directly allied to extant mammalian clades (Benton, 2004). Nevertheless,

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Figure 2. View along strike toward the south. From oldest on the right to youngest on the left: Early Jurassic Aztec Sandstone, Lower Cretaceous Willow Tank Formation, Upper Cretaceous Baseline Sandstone.

Figure 3. A mottled, massive mudrock with carbonate nodules. This horizon is interpreted as a paleosol horizon. Note, to left of scale bar is a root cast.

Figure 4. Single storied sandstone bodies circled in white. These are interpreted to be channel ﬁll deposits. View toward the south, oblique to strike.

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Early Cretaceous mammals show surprising diversity and include scansorial, semi-aquatic, arboreal, gliding and fossorial forms (Ji et al., 2006; Meng et al., 2006). In addition, cryptodiran turtles and squamates continue their Late Jurassic diversiﬁcation, and snakes make their ﬁrst appearance in the fossil record (Benton, 2004). Lagerstatten evidence also indicates a surprising diversity in pterosaur size and ecologic niches in the Early Cretaceous, which likely exceeded that of Late Cretaceous taxa (Wang and Zhou, 2006). Wang and his colleagues (2005b) suggest that pterosaurs predominated in coastal areas, whereas birds occupied terrestrial, inland regions during the Early Cretaceous. From the single Jurassic genus, Archaeopteryx, birds diversify into a variety of forms and clades (Chiappe, 2007). Although current fossil evidence provides inadequate evidence for the early diversiﬁcation of modern bird orders, as predicted by molecular clock data, ample evidence supports a major avian radiation in Early Cretaceous time. Several lineages of nonavian dinosaurs characterize but are not restricted to the Lower Cretaceous (e.g., Spinosauridae and Carcharodontosaurididae) (Holtz et al., 2004). Most Lower Cretaceous theropods include Jurassic holdovers such as a few ceratosaurs and compsognathids, as well as a variety of basal tetanurans (Tykoski and Rowe, 2004; Holtz et al., 2004). Tyrannosauroidea and the Abelisauridae are important Late Cretaceaous theropod clades but are only sparsely represented in the Lower Cretaceous (Holtz, 2004; Sereno et al., 2004; Xu et al., 2004). More derived tetanurans, troodontids, dromaeosaurids, oviraptorids, and therizinosaurids are also limited in their occurrence and restricted to Asia and/or North America (Clark et al., 2004; Makovicky and Norell, 2004; Norell and Makovicky, 2004; Osmolska et al., 2004). Representative taxa from North America include the dromaeosurids Deinonychus and Utahraptor, the oviraptorid Microvenator, and the therizinosaurid Falcarius. Sauropods remained an important component of terrestrial faunas worldwide throughout the Early Cretaceous. Principal taxa were Diplodocoidea and the Macronaria (Upchurch et al., 2004); the latter includes basal titanosauriforms, brachiosaurids, and titanosaurs. North American formations record Pleurocoelus in the Arundal Formation of Maryland, Sauroposeidon in the Antlers Formation of Oklahoma, and Cedarusaurus and Venenosaurus from the Cedar Mountain Formation of Utah (Upchurch et al., 2004). Recently, an unidentiﬁed titanosaurid, represented by a single tooth, was recovered from the Willow Tank Formation. Ornithischian dinosaurs of the Lower Cretaceous primarily consist of ankylosaurians and ornithopods. Stegosaur diversity had largely waned and marginocephalians had yet to diversify signiﬁcantly (Galton and Upchurch, 2004; Maryanska et al., 2004; Hailu and Dodson, 2004). Marginocephalians have a limited geographical and taxonomic distribution in the Lower Cretaceous, whereas among ornithopods, hypsilophodonts constitute an important faunal component in some regions such as Europe and high-latitude Australia assemblages (Norman et al., 2004). In the Lower Cretaceous, more derived iguanodontians are abundant and common worldwide, with the exception of South America (Norman, 2004). For example, Tenontosaurus represents the principal

herbivore in Aptian/Albian Cloverly Formations of Montana and Planicoxa and Eolambia occur in the Cedar Mountain Formation of Utah. Iguanodons eventually yield to more derived hadrosaurid ornithopods in the Late Cretaceous (Norman, 2004). The Early Cretaceous spans 45 million years. This extended period of time, combined with the fragmentation of Laurasia and Gondwana led to a diversity of terrestrial faunas. Although difﬁcult to categorize, there are some unifying trends: Spinosaurs, carcharodontosaurids, iguanodons, rebbachisaurid sauropods, and nodosaurids are the most characteristic dinosaur lineages of the Early Cretaceous. Monotreme, marsupial, and placental mammals, a variety of primitive avian groups, as well as snakes originate in and continue to diversify through this time interval. In general, the changing environments of the Early Cretaceous are reﬂected in the differentiation of the faunas of this time. PALEONTOLOGY OF THE WILLOW TANK FORMATION Fossil material collected from 10 localities in the Willow Tank Formation produced a total of 16 taxa, including invertebrates, plants, dinosaurs, crocodilians, turtles, and ﬁsh (Tables 1 and 2; Figs. 5 and 6). These taxa are nearly evenly distributed among terrestrial and semi-aquatic to aquatic groups. Single elements comprise most of the vertebrate material, with the dinosaurs represented primarily by teeth. The dinosaur assemblage consists of dromaeosaurid, tyrannosauroid, titanosauriform, iguanodontian, and thyreophoran taxa. Only two dinosaur localities yielded associated osteological remains; these sites contain dromaeosaur and ornithopod elements. Additional localities produced dinosaur eggshell (discussed below). The only bonebed discovered thus far consists of turtle remains preserved in ﬁne-grained overbank sediments. The bonebed contains at least three, possibly four species of turtles and includes at least three species of ﬁsh, represented by two types of ganoid scales and two dipnoan tooth plates. Only a single trace fossil, Scoyenia, has been recognized in these poorly indurated lithologies. Only one gastropod shell documents the presence of freshwater invertebrates. Six of the ten fossil-bearing localities in the upper member of the Willow Tank formation represent well-drained ﬂoodplain environments; these localities include microsite accumulations or sites with isolated elements. One sandstone, interpreted as an avulsed channel ﬁll, produced vertebrate remains. In contrast to these sites, the fern locality occurred in a volcanogenic deposit. The other two localities consisted of ﬂoat material and the depositional environment was indeterminate. Early Cretaceous faunas across North America appear rather homogenous at the family level, with relatively consistent distribution of turtle and dinosaur taxa across the continent (Gilmore, 1922; Ostrom, 1970; Winkler et al., 1990; Kirkland et al., 1997; Kirkland et al., 1998). The taxa present within the Willow Tank Formation are similar to those found in other Lower Cretaceous formations from North America. For example, the one turtle genus present in Willow Tank Formation, Naomichelys, is also

Dinosaurs and dunes TABLE 1. LIST OF WILLOW TANK FORMATION VERTEBRATE TAXA Vertebrates Dinosauria Theropoda Dromaeosauridae Tyrannosauridae Indeterminate Family Ootaxa Theropod type I Theropod type II Sauropoda Titanosauriformes Ornithopoda Iguanodontia plus an advanced Iguanodontian? Thyreophora Crocodilia Chelonia Cryptodira Baenidae Naomichelys c.f. Adocus ?Trioncychidae? Osteichthys Ceratodus Lepisosteidae Holostean A

TABLE 2. LIST OF NON-VERTEBRATE WILLOW TANK FORMATION TAXA Invertebrate Mollusca Gastropoda Plantae Pterphyta Morphotype 1 Morphotype 2 Ichnotaxon Coprolites Scoyenia

found in the majority of Lower Cretaceous formations of North America. Another common Lower Cretaceous turtle genus, Glyptops, and Naomichelys occur in the Arundel Formation of Maryland (Hay, 1908; Kranz, 1998), the Cloverly Formation of Montana and Wyoming (Ostrom, 1970), the Cedar Mountain Formation of Utah (Kirkland et al., 1997; Cifelli et al., 1999), and the Trinity Group of the southern coastal plain (Nydam et al., 1997). The fauna of the Willow Tank Formation is most similar to those present in the Cedar Mountain Formation of Utah and the Cloverly Formation of Wyoming and Montana. This similarity potentially expands the biogeographic range of this fauna to the southwest and the margin of the fold and thrust front. THE FIRST FOSSIL EGGSHELLS FROM NEVADA In addition to the vertebrate osteological remains, the Willow Tank Formation recently produced the ﬁrst dinosaur eggshells

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from Nevada. Cretaceous eggs and eggshells from deposits older than the Campanian in North America are rare. The Cedar Mountain, Dakota, and Kelvin formations of Utah provide the best-documented Early Cretaceous fossil eggshell localities in the southwestern United States (Jensen, 1970; Bray, 1998; Carpenter, 1999; Zelenitsky et al., 2000). Although intact or even partial eggs have yet to be discovered in Valley of Fire State Park, several sites produced eggshell fragments, including a microsite and two localities that occur in ﬁne-grained overbank deposits. The eggshells from these localities represent two types of avian or non-avian theropod eggs. The ﬁrst type of theropod eggshell (VF 3-1) measures 1.6– 2.0 mm and consists of two structural layers of calcite (Fig. 7A). The eggshell is similar to Macroelongatoolithus carylei (oofamily Elongatoolithidae), an oospecies known only from eggshell fragments from the Lower Cretaceous Cedar Mountain and Kelvin formations of Utah. Elongate eggs from Asia, referable to this oogenus, sometimes exceed 50 cm in length and are thought to represent a very large theropod (Jin et al., 2007). However, the proportions of the two structural layers in M. carylei and the Valley of Fire eggshell differ, thus suggesting the latter may represent a different oospecies. The second type of eggshell (VF 2-3_4) measures 0.6– 0.8 mm thick, exhibits three structural layers of calcite, and displays prominent nodes on the outer shell surface (Fig. 7B; Table 3). From the inner to the outer eggshell surfaces, the layers include the mammillary layer, cryptoprismatic layer, and external layer. The presence of an external layer in the Valley of Fire eggshell distinguishes this eggshell from all oospecies in the Elongatoolithidae (Fig. 7B). Although an external layer is sometimes reported in Cenozoic eggshells (Mikhailov, 1997; see also Kohring, 1999 and references therein), this structural attribute is relatively rare in Mesozoic eggs and occurs primarily in small eggs with thin shells. These small eggs are similar in thickness to modern neognathid bird eggs (i.e., 0.2–0.3 mm) and lack surface ornamentation. The relatively thicker shell of the Valley of Fire specimen in comparison to other eggs currently assigned to the oofamily Neognathoolithidae, suggests the Nevada specimen may represent a new egg type. FIELD TRIP ROAD LOG AND DESCRIPTION OF STOPS The ﬁeld trip begins at the University of Nevada–Las Vegas campus. Drive north 37 miles on I-15 to Exit 75, the Valley of Fire and Lake Mead exit. Continue east ~13 miles to the entrance of the Valley of Fire State Park. Within the park, somber graygreen outcrops on the right (south) that parallel the highway are Paleozoic in age. Isolated remnants of the bright red Jurassic Sandstone occasionally crop out between the Paleozoic section and the highway. Approximately 2 miles past the park entrance turn left at the sign for the Atlatl Rock campground. Stay to the right past the turnoff to the campground, then turn left into the Atlatl Rock picnic area.

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Figure 5 (on this and previous page). Taxa of the Willow Tank Formation. (A) Ceratodus tooth. (B) Crocodilian scute. (C) Adocus costal fragment. (D) Baenid costal fragment. (E) Naomichelys costal fragment. (F) Iguanodontid tooth. (G) Titanosauriform tooth. (H) Lateral view of Tyrannosauroid tooth. (I) Occlusial view of same Tyrannosauroid tooth.

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Figure 6. (A) Dromaeosaur femur. (B) Thyreophoran tooth. (C) Iguanodontian pubis. (D) Fern morphotype 1. (E) Fern morphotype 2.

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Stop 1: Atlatl Rock—Jurassic Aztec Sandstone Atlatl Rock is one of the most popular tourist attractions in the Valley of Fire State Park. A metal staircase provides access to petroglyphs high on the cliff of the Jurassic Aztec Sandstone. The petroglyphs inscribed into the dark-stained desert varnish covering the rock face depict ﬁgures and enigmatic symbols of Native American peoples who inhabited the area ~3000 years ago. One such ﬁgure is the atlatl, a notched stick used for spear throwing and a precursor to the bow and arrow. The desert varnish consists primarily of clay particles that catch additional substances that chemically react together when the rock reaches high temperatures in the desert sun. Iron and unusually high concentrations of manganese, thought to result from microbial activity, oxidize to form the dark patina that stains the sandstone. Desert varnish provided a natural easel for early Native American people to record their history, as indicated by the many petroglyphs present in the area of Atlatl Rock. The Aztec Sandstone overhangs the petroglyphs at the top of the staircase, providing a view of mammal-like reptile tracks on the bottom of the exposed bedding surface. This series of approximately eight prints (preserved as hyporelief) are similar to betterpreserved tracks at another locality in the park. Rowland and Mercadante (2007) attributed these trace fossil tracks to Brasilichnium; this ichnotaxon also occurs in exposures of the Aztec Sandstone in eastern California. The tracks recently discovered in the vicinity of Atlatl Rock differ from Brasilichnium in that they measure only ~2 cm long. These tracks, however, have yet to be studied. Exit Atlatl Rock picnic area by the same route, returning to the main road. Turn left and drive ~2 miles east to the turnoff to the Visitor’s Center on the left. Continue past the Visitor’s Center up the steep hill for ~2 miles, stopping at the Rainbow Vista parking lot on the right side of the road. Stop 2: Rainbow Vista—Jurassic Aztec Sandstone

Figure 7. Theropod eggshells from Valley of Fire State Park. (A) Type 1 eggshell showing two structural layers of calcite. (B) Type 2 eggshell showing three structural layers. (C) Same eggshell as B, showing wide pore (P). Layer 1 is absent due to weathering. Scale bars equal 100 µm.

Rainbow Vista provides spectacular views of the Lower Jurassic Aztec Sandstone that dominates Valley of Fire State Park (Fig. 1). The multicolored alteration patterns in the rocks reﬂect syndepositional reddening of the eolian sandstone and repeated episodes of dissolution, mobilization, and reprecipitation of iron oxide and hydroxide (Eichhubl et al., 2004). Due to structural deformation of this area, Cretaceous-age rocks occur only east of the main N-S route through Valley of Fire. To the northeast, outcrops of the Willow Tank Formation are identiﬁable by their gray banding, which contrast to the bright red of the Aztec Sandstone; however, exposures visible from this location are limited. From the parking lot at Stop 2 turn right onto the main road and drive ~2.5 miles north to the end of the White Domes turnabout. White Domes picnic area includes tables and restroom accommodations. After a half-hour lunch, drive south on the main road, toward the Visitor’s Center. After traveling a short distance (~1.5 miles), park at the second turnoff past White Domes called Parking Lot #3.

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Bonde et al. TABLE 3. COMPARISON OF MACROELONGATOOLITHUS CARLYLEI TO TWO THEROPOD EGGSHELLS FROM VALLEY OF FIRE STATE PARK, NEVADA Macroelongatoolithus carlylei Valley of Fire Theropod Valley of Fire (Zelenitsky et al., 2000) Type I Theropod Type II Egg size unknown Unknown unknown Number of Layers 2 2 3 Shell Thickness 1.38–3.0 1.6–2.0 0.75 Layer 1/Layer 2 ratio 1:2 to 1:4 1:6 1:4 to l:5.5 Note: All measurements in mm.

Stop 3: Parking Lot #3—The Willow Tank Formation From the parking lot we will cross the road and go downhill until we meet a major wash. Enter the major wash and go to the east and walk about one-half mile. The route crosses extensive exposures of the Aztec Sandstone and passes an abrupt contact between the Aztec Sandstone and the basal conglomerate member of the Willow Tank Formation. From this main wash, enter the tributary wash to the north that cuts into the easily weathered, mudrocks of the Willow Tank Formation. The fossiliferous exposures of interest are approximately one-third of a mile to the north. At this locality, a grayish mudrock overlies a laterally continuous white tephra layer. Interpreted as an overbank deposit, this mudrock produced most of the turtle material recovered from the Willow Tank Formation and represents the only bonebed yet discovered. Additional exposures of the upper member of the Willow Tank Formation are located nearby and preserve a carbonate nodule horizon and isolated, lenticular sandstone lenses, typical of the formation. These lenticular sandstone lenses are interpreted to be channel ﬁll deposits from the anastomosed ﬂuvial system. Dinosaur-bearing units are limited to this upper member of the Willow Tank Formation. Continuing ~100 m to the north, enter the wash to the west, marked by boulders of the basal conglomerate. The route leads through a small, photogenic slot canyon cut into the Aztec Sandstone and eventually returns to the main road. Turn left and walk up the steep hill to the parking lot and vehicles. If time permits, we will stop at the Visitor’s Center, before returning to Las Vegas. The park facility provides information and exhibits about the ﬂora and fauna of the park, as well as the Native Americans who once inhabited the area. REFERENCES CITED Allen, J.R.L., 1977, Physical processes of sedimentation: An introduction: London, UK, George Allen and Unwin Ltd., 4th edition, 248 p. Archer, M., Flannery, T.F., Ritchie, A., and Molnar, R.E., 1985, First Mesozoic mammal from Australia—An Early Cretaceous monotreme: Nature, v. 318, p. 363–366, doi: 10.1038/318363a0. Ash, A.R., and Read, C.B., 1976, North American species of Tempskya and their stratigraphic signiﬁcance: U.S. Geological Survey Professional Paper 874, 42 p. Aydin, A., 1978, Small faults formed as deformation bands in sandstone: Pure and Applied Geophysics, v. 116, p. 913–930, doi: 10.1007/BF00876546. Barron, E.S., 1983, A warm, equable Cretaceous; the nature of the problem: Earth-Science Reviews, v. 19, no. 4, p. 305–338, doi: 10.1016/0012-8252 (83)90001-6. Benton, M.J., 2004, Vertebrate Palaeontology: Malden, Massachusetts, Blackwell Publishing, p. 455.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 28 JANUARY 2008

Printed in the USA

Contents Preface 1. The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section

Denise Honn and Eugene I. Smith 2 Late Paleozoic deformation in central and southern Nevada

Pat Cashman, Jim Trexler, Walt Snyder, Vladimir Davydov, and Wanda Taylor 3 Active tectonics of the eastern California shear zone

Kurt L. Frankel, Allen F. Glazner, Eric Kirby, Francis C. Monastero, Michael D. Strane~ Michael E. Qskin-' .jeffrey R. Unruh, J. Douglas Walker, Sridhar Anandakrishnan, John M~ Bartley, Drew~· r;oleman.L .james F. Dolan.L Robert Finke~ Dave Greene.L Andrew Kylander-Ciark.L Shasta Marrero-' Lewis A. Qwen.L and Fred Phillips

r

4. Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic-Paleozoic boundary

Stephen M. Rowland, Lynn K. Oliver, and Melissa Hicks 5. Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert

Andrew P. Barth, J. Lawford Anderson, Carl E. Jacobson, Scott R. Paterson, and Joseph L. Wooden

6. Cenozoic evolution ofthe abrupt Colorado Plateau-Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River

James E. Faulds, Keith A. Howard, and Ernest M. Duebendorfer 7. Interpretation of Pleistocene glaciation in the Spring Mountains of Nevada· Pros and Cons

Jerry Osborn, Matthew Lachniet, and Marvin (Nick) Saines 8. Quaternary volcanism in the San Francisco Volcanic Field: Recent basaltic eruptions that profoundly impacted the northern Arizona landscape and disrupted the lives of nearby residents

S.L. Hanson, W Duffield, and J. Plescia 9. The ~rit Mountain batholith and Secret Pass Canyon volcanic center: A cross-sectional view of the magmatic architecture of the uppermost crust of an extensional terrain, Colorado River, Nevada-Arizona

Nicholas P. Lang, B.J. Walker, Lily L. Claiborne, Calvin F. Miller, Richard W Hazlett, and Matthew T. Heizler 10. Devonian carbonate platform of eastern Nevada: Facies, surfaces, cycles, sequences, reefs, and catastrophic Alamo Impact Breccia

John E. Warme, Jared R. Morrow, and Charles A. Sandberg 11. Dinosaurs and dunes! Sedimentology and paleontology of the Mesozoic in the Valley of Fire State Park

Joshua W Bonde, David J. Varricchio, Frankie D. Jackson, David B. Loope, and AubreyM. Shirk

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