Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera (GSA Field Guide 21)

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Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera

edited by Jeffrey Lee Department of Geological Sciences 400 E. University Way Central Washington University Ellensburg, Washington 98926 USA James P. Evans Department of Geology Utah State University 4505 Old Main Hill Logan, Utah 84322-4505 USA

Field Guide 21 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140, USA

2011

Copyright © 2011, The Geological Society of America (GSA), Inc. 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. In addition, an author has the right to use his or her article or a portion of the article in a thesis or dissertation without requesting permission from GSA, provided the bibliographic citation and the GSA copyright credit line are given on the appropriate pages. 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. GSA provides this and other forums for the presentation of diverse opinions and positions by scientists worldwide, regardless of their race, citizenship, gender, religion, sexual orientation, or political viewpoint. Opinions presented in this publication do not reflect official positions of the Society. 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. Cataloging-in-Publication Data for this volume is available from the Library of Congress. Cover: Cedar Creek and Borah Peak horst. Borah Peak fault scarp is plain at the base of the mountains. Cedar Creek contains multiple generations of late Pleistocene moraines, cut by the Borah Peak fault. See Figure 18 in Chapter 5 (P.K. Link and M.K.V. Hodges, The Neogene drainage history of south-central Idaho, p. 103–123).

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Contents

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. New investigations of Pleistocene glacial and pluvial records in northeastern Nevada . . . . . . . . . 1 Jeffrey S. Munroe and Benjamin J.C. Laabs 2. Timing, distribution, amount, and style of Cenozoic extension in the northern Great Basin . . . . 27 Christopher D. Henry, Allen J. McGrew, Joseph P. Colgan, Arthur W. Snoke, and Matthew E. Brueseke 3. Tectonomagmatic evolution of distinct arc terranes in the Blue Mountains Province, Oregon and Idaho . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 C.J. Northrup, M. Schmitz, G. Kurz, and K. Tumpane 4. Neogene drainage development of Marsh and Portneuf valleys, eastern Idaho . . . . . . . . . . . . . . 89 Glenn D. Thackray, David W. Rodgers, and Andrew Drabick 5. The Neogene drainage history of south-central Idaho . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 Paul K. Link and Mary K.V. Hodges 6. Paleontology and stratigraphy of middle Eocene rock units in the Bridger and Uinta Basins, Wyoming and Utah . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 Paul C. Murphey, K.E. Beth Townsend, Anthony R. Friscia, and Emmett Evanoff 7. Middle Cryogenian (“Sturtian”) Pocatello Formation: Field relations on Oxford Mountain and the Portneuf area, southeast Idaho . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167 Joshua A. Keeley and Paul K. Link 8. New descriptions of the cap dolostone and associated strata, Neoproterozoic Pocatello Formation, southeastern Idaho, USA . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 183 Carol M. Dehler, Kathleen Anderson, and Robin Nagy 9. Reinterpreted history of latest Pleistocene Lake Bonneville: Geologic setting of threshold failure, Bonneville flood, deltas of the Bear River, and outlets for two Provo shorelines, southeastern Idaho, USA . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 195 Susanne U. Janecke and Robert Q. Oaks Jr.

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Preface

The geologic record preserved within the Rocky Mountain and Cordilleran regions spans nearly all of Earth’s time. The combination of the long geologic record and stunning scenery has attracted geologists for two centuries. Past and ongoing geologic research in this region has resulted in a wealth of significant observations and paradigm shifts in interpretations. This field trip guide, compiled for the 2011 joint meeting of the GSA Rocky Mountain and Cordilleran Sections, provides a small and succulent appetizer to the full menu of remarkable geology of the Rocky Mountain and Cordillera regions. Field trips presented in this volume span geologic topics from Neoproterozoic deposits, late Paleozoic–early Mesozoic terrane accretion, Eocene mammals and climate, Eocene to middle Miocene extension, late Miocene and younger basin and river system evolution, and Pleistocene glaciers and pluvial lakes. In chapter one, Munroe and Laabs’ field trip examines latest Pleistocene to Holocene mountain glacial and pluvial lake deposits in the East Humboldt and Ruby Mountains, northeastern Nevada, and implications for temperature and precipitation changes during this time period. Henry et al.’s field trip, in chapter two, highlights geologic field relations, quantitative metamorphic petrology, and extensive geochronology that indicate a contrast in the timing of mid-crustal versus supracrustal extension in the East Humboldt and Ruby Mountains area, northeastern Nevada. In chapter three, Northrup and colleagues discuss new mapping, geochronology, and geochemistry data that bear on the tectonic evolution and pre-accretion history of the Wallowa and Olds Ferry terranes of the Blue Mountains Province, Oregon and Idaho. Thackray et al., in chapter four, investigate the basin evolution of Marsh and Portneuf valleys, Idaho. The focus of this trip is to examine the basin’s complex record of drainage capture, faulting, volcanism, and the Bonneville Flood. In chapter five, a companion paper to Thackray et al., Link and Hodges summarize detrital-zircon geochronology from two large river systems—the Big Lost and Wood rivers—on the north side of the Snake River Plain that shed light on the course of these rivers during the past 10 m.y. and the location of the paleo-continental divide. Murphey et al.’s field trip, described in chapter six, centers on a 10-m.y.-long biotic, environmental, and climate history recorded in Eocene sediments of the Green River and Uinta basins, Utah. The next two chapters are for a trip focusing on the Neoproterozoic Pocatello Formation, Idaho. Keeley and Link, in chapter 7, examine the diamictite stratigraphy and discuss U-Pb zircon ages from these deposits in the Oxford Mountain and Portneuf area. In chapter eight, Dehler et al. investigate the stratigraphy and sedimentary facies of deposits that overlie diamictites in the Fort Hall Mine area south of Portneuf Narrows. In chapter nine, Janecke and Oaks highlight geologic, geomorphic, and geophysical data that document the southward shift of the Lake Bonneville outlet. Thanks to the authors for agreeing to lead a field trip, to reviewers for their thoughtful and helpful reviews, and the GSA publications staff for their time and effort in compiling this volume. Jeffrey Lee James Evans

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The Geological Society of America Field Guide 21 2011

New investigations of Pleistocene glacial and pluvial records in northeastern Nevada Jeffrey S. Munroe Geology Department, Middlebury College, Middlebury, Vermont 05753, USA Benjamin J.C. Laabs Department of Geological Sciences, State University of New York–Geneseo, Geneseo, New York 14454, USA

ABSTRACT The Great Basin of the western United States offers tremendous potential for exploring the response of mountain glaciers and lowland lakes to climate changes during the Last Glacial Maximum (LGM, MIS-2, ca. 22–18 ka B.P.) and subsequent glacial-interglacial transition. The combination of well-distributed alpine moraine records and pluvial lake deposits offers an unparalleled opportunity to develop a more precise understanding of temperature and precipitation changes during the latest Pleistocene and into the Holocene. This field trip provides an overview of recent and ongoing work illuminating aspects of the glacial and pluvial history of northeastern Nevada from the LGM to the present. The route of this trip involves three full days of stops separated by two nights in Elko, Nevada. The first day focuses on glacial deposits at the type locality for the LGM Angel Lake Glaciation on the eastern side of the East Humboldt Range. The second day explores the geomorphic record of pluvial Lakes Franklin and Clover on the east side of the Ruby–East Humboldt Mountains and describes recent efforts to develop a chronology for the late Pleistocene regression of these lakes. The final day again focuses on glacial geology, starting with the type locality of the pre-LGM Lamoille Glaciation on the west side of the Ruby Mountains, and ending with several stops along the scenic drive up Lamoille Canyon.

INTRODUCTION

2001; Blackwelder, 1934). For instance, the Sierra Nevada, which form the western border of the Great Basin, were extensively glaciated and were one of the first locations in which cosmogenic surface-exposure dating was employed to develop a moraine chronology (Phillips et al., 1996). At the eastern border of the Great Basin, a detailed chronology of latest Pleistocene glaciation has been developed for the Wasatch Mountains (Madsen and Currey, 1979; Lips et al., 2005; Laabs et al., 2007) and glacial records there have been used in numerical modeling exercises

The Great Basin of the western United States offers tremendous potential for exploring the response of mountain glaciers and lakes to climate changes during the Last Glacial Maximum (LGM) and subsequent glacial-interglacial transition (GIT). Despite the modern arid climate and hot summer temperatures, numerous mountain ranges in the region contain evidence of former glaciers (Blackwelder, 1931; Sharp, 1938; Osborn and Bevis,

Munroe, J.S., and Laabs, B.J.C., 2011, New investigations of Pleistocene glacial and pluvial records in northeastern Nevada, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 1–25, doi: 10.1130/2011.0021(01). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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to yield paleoclimatic inferences (Plummer and Phillips, 2003; Laabs et al., 2006). Yet in contrast to these well-studied ranges at the margins, little is known about the timing of late Pleistocene glacier fluctuations in the interior ranges of the Great Basin, an area covering more than 500,000 km2. The largest glaciers between the Wasatch and Sierra Nevada were located in the East Humboldt and Ruby Mountains of northeastern Nevada where moraines document the LGM and multiple glacial stillstands during the last GIT (Laabs and Munroe, 2008). Indeed, the last glaciation in the Great Basin was designated the “Angel Lake Glaciation” after well-preserved glacial deposits in the East Humboldt Range (Sharp, 1938). Three pluvial lakes were also present near the East Humboldt and Ruby Mountains during the Angel Lake Glaciation: Lakes Franklin, Clover, and Waring. Although these lakes were small compared with the better-known Lakes Lahontan and Bonneville, their deposits are easily recognized and mapped, allowing their former extents to be delineated (Reheis, 1999). The combination of well-distributed alpine moraine records and pluvial lake deposits offers an unparalleled opportunity to develop a more precise understanding of temperature and precipitation changes during the latest Pleistocene and into the Holocene. TRIP OVERVIEW This field trip provides an overview of recent and ongoing work illuminating aspects of the glacial and pluvial history of northeastern Nevada from the Last Glacial Maximum (MIS-2, ca. 22–18 ka B.P.) to the present. The route of this trip involves

three full days of stops separated by two nights in Elko, Nevada (Fig. 1). Latitude and longitude coordinates are given for each stop (WGS-84) in the following descriptions as an aid to future users of this guide. The first day of the trip focuses on glacial deposits at the type locality for the LGM Angel Lake Glaciation on the eastern side of the East Humboldt Range. From Logan, we will drive west across the Bonneville Basin in northwestern Utah into Nevada. After reaching I-80, we will continue west over Pequop Summit, across the Independence Valley, and over Moor Summit to the town of Wells. Exiting the highway at Wells, we will follow a winding road up to the valley of Angel Creek and spend the afternoon visiting three closely spaced stops (Fig. 2). After a night in Elko, Nevada, the second day of the trip explores the geomorphic record of pluvial Lakes Franklin and Clover on the east side of the Ruby–East Humboldt Mountains (Fig. 2). The Franklin Valley is home to the Ruby Lake National Wildlife Refuge, and we will start our day with an overview of the Refuge provided by staff of the U.S. Fish and Wildlife Service. We will then work our way northward from the southern end of the Franklin Valley, stopping to examine localities where the late Pleistocene highstand and other water planes are well expressed. We will also discuss new geochronologic constraints on some of these deposits. In mid-afternoon, we will exit the north end of the former Lake Franklin basin and cross the divide to the Clover Valley that hosted pluvial Lake Clover during the late Pleistocene. Our two final stops will provide an overview of the Lake Clover geomorphic record, including a spectacular series of beach ridges that have been dated by optically stimulated luminescence (OSL) and 14C.

Figure 1. Route of the field trip from Logan, Utah (UT), to northeast Nevada (NV). F—Lake Franklin; C—Lake Clover; W—Lake Waring. Lower right inset shows location of figure in western United States. Shaded box outlines the location of Figure 2.

Pleistocene glacial and pluvial records in northeastern Nevada The final day will again focus on glacial geology, starting with the type locality of the pre-LGM Lamoille Glaciation on the west side of the Ruby Mountains (Fig. 2). Deposits at the mouth of Lamoille Canyon have been assumed to represent an advance during MIS-6, and we will discuss efforts to date these landforms with cosmogenic 10Be surface-exposure dating. We will also hike southward along the range front to Seitz Canyon to visit a spectacularly preserved series of Lamoille and Angel Lake–age moraines that have been the focus of a concerted 10Be dating effort. Our final stops will be along the scenic drive up Lamoille Canyon, a deep U-shaped valley considered the “Yosemite of Nevada.” We will discuss the geomorphology of the valley, the distribution of glacial deposits within it, and efforts to identify the moraine from the LGM (i.e., the Angel Lake equivalent moraine in Lamoille Canyon). After lunch at the head of Lamoille Canyon, we will drive northward along the western slope of the Ruby–East Humboldt Mountains to reach I-80 and retrace our route back to Logan.

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PHYSICAL SETTING Northeastern Nevada is the heart of the Great Basin and features classic Basin and Range topography. The relative lowland of the Bonneville Basin, formerly inundated by pluvial Lake Bonneville, ends abruptly at the Utah-Nevada state line, where the landscape takes on an increasingly corrugated appearance moving westward from the border. Long, linear valleys separated by rugged north-south–oriented mountain ranges characterize the area of this field trip. Moving westward from the state line along I-80, one sequentially crosses the Toano Range at Silverzone Pass (1815 m), the Goshute Valley formerly occupied by pluvial Lake Waring, the Pequop Mountains at Pequop Summit (2124 m), the Independence Valley, Moor Summit (1882 m), and the Clover Valley before arriving at the highest mountains in the region, the Ruby– East Humboldt Range. These mountains extend for ~130 km with an orientation of north-northeast to south-southwest. Summit elevations range up to 3471 m at Ruby Dome.

Figure 2. Detail map of field trip stops. Dotted line shows alternate route over Harrison Pass on Day 2. See text for details.

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The modern drainage systems in this area are ephemeral and primarily carry spring snowmelt and precipitation from summer thunderstorms down to evaporate on basin floors. The northwestern end of the Ruby–East Humboldt Range, however, is within the headwaters of the Humboldt River which flows for ~500 km westward across northern Nevada before terminating in the Humboldt Sink. Interstate 80 follows the Humboldt River from near Wells, downstream to the Sink in a broad arc across northern Nevada; however, high mountains with obvious glacial geomorphology are generally absent in this sector of the state. GLACIAL GEOMORPHOLOGY Although glaciers are absent from the mountains of northeastern Nevada today, the spectacular upland landscapes, particularly of the Ruby–East Humboldt Range, testify to extensive Pleistocene glaciation. Glaciers formed at the heads of most valleys draining the mountain flanks. Some of these were short and remained confined to their cirques. Others flowed at least some distance downvalley, converting valley cross sections into broader U-shaped profiles. Given the asymmetric nature of the Ruby–East Humboldt uplift, valleys are much steeper on the eastern mountain flank. Glaciers in some of these valleys must have resembled icefalls more than typical valley glaciers. In contrast, glaciers descending the gentler western valley slopes flowed a greater distance from cirque headwalls, but reached similar terminus elevations. Blackwelder (1931; 1934) presented a short overview of the glacial geology of the Ruby Mountains. In colorful language, he described the “wild crags of the freshly torn cirques” and noted that the glacial deposits represent two glaciations separated by a prolonged interval of weathering. He named these the Lamoille (older) and Angel Lake (younger) glaciations. Sharp (1938) expanded upon Blackwelder’s work, formalized the Lamoille and Angel Lake designations, discussed controls over the extent of glaciation, and summarized mechanisms of post-glacial valley modification. Terminal moraines dating to the Angel Lake Glaciation form bouldery arcuate ridges that cross the floors of most valleys. Streams have cut narrow post-glacial channels through these, but the moraine forms are still obvious. In many settings, multiple crests are present on the overall Angel Lake moraine complex. Older moraine ridges representing the Lamoille Glaciation are found at lower elevations in most valleys. At the type locality at the mouth of Lamoille Canyon, these deposits delineate a piedmont lobe that sprawled outward from the mountain front. In other valleys, particularly along the western slope, Lamoille deposits form steep-sided lateral moraines that project beyond the Angel Lake terminal moraine complex. Other than at Lamoille Canyon, terminal moraines of Lamoille age are not preserved. Downslope from these features, outwash valley trains and meltwater channels are locally well developed. Recent inventorying of glacial features has supported a comprehensive reconstruction of former glaciers in the Ruby–East

Humboldt Mountains. This work indicates that the Ruby–East Humboldt Mountains contained over 130 glaciers during the Angel Lake Glaciation, which collectively covered more than 270 km2 (Laabs and Munroe, 2008). Reconstructed equilibrium line altitudes (ELAs) range from 2350 to 3000 m, with a mean of ~2700 m. Equilibrium lines generally lowered in elevation from south to north, and 10 of the 15 glaciers with ELAs below 2500 m were located on the western slope. Together, this pattern indicates prevailing moisture transport from the northwest during the Angel Lake Glaciation. In a comprehensive overview of glacial deposits in the Great Basin, it was noted that Angel Lake moraines are commonly bulky masses of till on valley floors that strongly contrast with the discrete ridges of the remaining Lamoille-age moraines (Osborn and Bevis, 2001). Such voluminous, hummocky moraines are usually produced by glaciers carrying large volumes of sediment. Because Angel Lake moraines typically exhibit this morphology throughout the Great Basin regardless of local lithology, there might be a climatic significance to this observation. Osborn and Bevis (2001) suggested that Angel Lake glaciers may have advanced, retreated, and readvanced multiple times to the same locations, eventually depositing large amounts of till in compound, terminal moraines. In support of this theory, they noted rock flour records from the Owens Valley which reveal multiple millennial-scale intervals of expanded ice in the Sierra Nevada leading up to the LGM (Benson et al., 1998; Bischoff and Cummins, 2001). An alternative possibility is that conditions in the Great Basin were too marginal for extensive alpine glaciation during MIS-4. Instead, this interval was marked by intense periglacial activity that generated large volumes of frost-shattered rubble in the valleys previously occupied by Lamoille glaciers. When the Angel Lake Glaciation began in MIS-2, glaciers transported this backlog of pre-weathered material downvalley to form the anomalously bulky terminal moraines. In contrast, other higher mountain regions in the western United States may have hosted glaciers during MIS-4 that cleaned their valleys down to bedrock. When alpine glaciers again formed and advanced in these valleys during MIS-2, only loose material generated during MIS-3 was available for immediate transport, leading to less voluminous moraines. PLUVIAL LAKES Despite the arid climate characterizing the valleys of northern Nevada today, abundant evidence exists for large lakes occupying valley floors in the past. These “pluvial” lakes have been the target of considerable study dating back to the pioneering work of prominent geologists with the U.S. Geological Survey (Gilbert, 1890; Russell, 1885). Features delimiting the former shorelines of these lakes are obvious in dozens of valleys (Reheis, 1999; Mifflin and Wheat, 1979). These include beach berms, cuspate spits, and lagoons at high elevations and featureless plains underlain by deepwater sediment at lower elevations. Mapping and dating of these landforms provides information about former

Pleistocene glacial and pluvial records in northeastern Nevada episodes of lake transgression and regression. Many of these lakes were hydrologically closed (i.e., lacking surface hydrologic connections to other basins), and thus fluctuations in former lake levels are important signals of past climate variability. The majority of work on pluvial lake records in the western United States has focused on the two largest lakes: Bonneville and Lahontan. At its maximum, Lake Bonneville covered ~50,000 km2 of western Utah and had a maximum depth of ~300 m (Oviatt, 1997). Lake Lahontan inundated ~22,000 km2 of interconnected basins in western Nevada and had a maximum depth of ~280 m (Mifflin and Wheat, 1979). The pre–MIS-2 history of both lakes is less precise, but available evidence indicates that lakes were present in these basins multiple times during the Quaternary (Oviatt et al., 1999). In contrast, the MIS-2 deposits of these lakes are well preserved, and abundant evidence indicates that both lakes rose to roughly coincident highstands during MIS-2. Regression continued throughout the latest Pleistocene, and the lakes were reduced to isolated playas and shallow, brackish ponds by the early Holocene (Oviatt et al., 2003). Despite their fame, Lakes Bonneville and Lahontan were not the only pluvial lakes in the Great Basin. Dozens of smaller lakes existed in valleys from New Mexico northwestward to Oregon, but the records from these settings are much less well understood (Reheis, 1999; Mifflin and Wheat, 1979). This trip in northeastern Nevada will pass through areas formerly inundated by pluvial Lakes Franklin, Clover, and Waring, the first two of which will be topics of numerous stops on Day 2. It is a logical starting assumption that these smaller pluvial lakes rose and fell in concert with Lake Bonneville and Lahonton. However, within this broad chronology, there are numerous details that have been (and remain) important topics for study. One topic is the relative timing of lake highstands across the region. The prevailing theory is that the pluvial climate responsible for the Great Basin lakes was a result of the southward deflection of prevailing storm tracks by the physical barrier imposed by the Laurentide Ice Sheet (Antevs, 1948). This mechanism, which has been broadly supported by numerical modeling (Bartlein et al., 1998), nicely explains the synchronicity between the global LGM and the rising pluvial lakes (Kutzbach, 1987). However, as the chronology of highstands in more basins has improved, a pattern has emerged in which southern lakes reached their highstands and began to regress while northern lakes were still rising (Garcia and Stokes, 2006; Enzel et al., 2003). This regional nonuniformity in lake behavior signifies important regional climatic variability during the last GIT. Most studies have attempted to explain this variability as a function of a northward migrating storm track following retreat of the Laurentide ice margin (Garcia and Stokes, 2006). However, additional work is clearly needed to determine the age of highstands in the numerous basins that remain undated. Nested within this growing recognition of regionally nonuniform lake behavior during the last GIT is geomorphic evidence for higher frequency climatic variability. Few basins contain just a single emergent shoreline; more commonly, multiple

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shorelines are preserved, reflecting multiple instances of balanced precipitation and evaporation that stabilized water levels and allowed formation of geomorphically conspicuous shoreline landforms during the overall regressive trend. What was the relative timing of these stillstands among basins? Were they driven by basin-specific factors, such as hypsometry, or do they contain a signal of climate variability during this crucial transition? We will consider these and other questions during our stops in the Franklin and Clover Valleys on Day 2. DAY 1. LOGAN TO ELKO: THE ANGEL LAKE TYPE LOCALITY (Total driving distance: ≈290 mi/467 km.) Day 1 Overview The first day of this field trip will focus on the region around Angel Lake in the East Humboldt Mountains near Wells, Nevada. Sharp (1938) designated Angel Lake as the type locality of the last Pleistocene glaciation in the Great Basin. Recent work has generated a more detailed view of the distribution of moraines in this valley and has attempted to assign ages to these landforms through cosmogenic 10Be surface-exposure dating. Paleolimnological investigations of sediment cores from Angel Lake have also shed light on Holocene environmental change in this area. We will drive to Angel Lake from Logan across the extreme northwest corner of Utah. After an early lunch upon arrival at the Angel Creek Campground, we will visit three stops at progressively higher elevations along the road leading up to Angel Lake. Please be prepared for hiking in rough and possibly wet terrain and for sudden changes in weather (temperatures at higher elevations may be close to 0 °C). Late in the afternoon we will continue west for another 45 minutes to Elko, Nevada, where we will spend the night at the Red Lion Hotel. Directions to Stop 1.1 From the Riverwoods Conference Center in Logan, begin driving west on E 200 N toward N Main Street. Continue on U.S.30 W toward I-15 S. Take exit 379 off I-15 to merge with I-84 W and travel 36.4 mi to exit 5. Continue on U.S.-30 across northwestern Utah, passing south of the Raft River Range and to the northwest of most of the former Lake Bonneville. At the UtahNevada border, U.S.-30 becomes NV-233. Note that Nevada is in the Pacific time zone; set clocks back 1 hour. Continue another 34.1 mi to I-80, and head west 27 mi to Wells, Nevada. Take exit 351, turn left under the highway, and then right on Angel Lake Road (NV-231). Follow this road 7.3 mi to the left turn into Angel Creek Campground (Fig. 3). Stop 1.1. Angel Lake Type Locality (41.02111°N, 115.08025°W) The drive from Wells, Nevada, toward Angel Lake affords many tremendous views of terminal moraines at the type locale

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for the Angel Lake Glaciation (Sharp, 1938). Here, in the northeastern sector of the East Humboldt Range, a 3-km-long glacier constructed latero-frontal moraines that form a nearly continuous terminal moraine loop in the Angel Creek drainage (Fig. 3). As noted by Osborn and Bevis (2001) the moraine crest (at 2317 m) is bouldery, displays low-relief hummocky topography, and has a steep ice-distal slope with ~60 m of relief (Fig. 4). Portions of the distal slope grade to an outwash fan that forms the surface at Stop 1.1. Elsewhere, till comprising the distal slope is inset to older, possibly Lamoille-equivalent moraines with relatively low-relief and few, deeply weathered erratic boulders at the crest. Along the southern slope of the terminal moraine, a recent breach in the moraine exposes till comprising the landform (Fig. 3). Although the exact origin of the breach is unknown, it apparently formed as a result of headward erosion in a tributary of the North Fork Angel Creek, perhaps aided by drainage of a small lake that may have existed behind the moraine loop. The exposed diamicton bears physical properties typical of sediments comprising alpine moraines; poorly sorted with a coarse-grained sandy matrix, clasts consisting of local lithologies (chiefly granodiorite and gneiss), and clast sizes ranging from pebbles to large boulders. Weathering profiles in the till are evident near the top of this and other exposures, and are consistent with those of

Angel Lake–equivalent till elsewhere in the region (Bevis, 1995; Wayne, 1984) and the correlation of the Angel Lake Glaciation to the Wisconsin Glaciation (Richmond, 1986). X-ray diffraction analysis of samples from within the weathering profile reveals the presence of pedogenic kaolinite and vermiculite that are absent in the unweathered till (Rosenberg et al., 2011). The precise timing of the Angel Lake Glaciation is poorly known throughout much of the Great Basin, due in large part to the lack of numerical age limits on glacial deposits. We are attempting to resolve this issue by applying cosmogenic 10Be surface-exposure dating to the type Angel Lake terminal moraine at this stop and to several other moraines in northeastern Nevada. As noted by Osborn and Bevis (2001), Angel Lake–equivalent moraines typically display bouldery crests; however, in the East Humboldt and Ruby Ranges, many erratic boulders are not suitable for cosmogenic 10Be surface-exposure dating because of their lithology (lacking quartz) or their degree of physical weathering. However, through extensive mapping we have identified numerous erratic boulders suitable for this dating method here and on other moraines in the Ruby and East Humboldt Ranges. Five erratic boulders atop the terminal moraine were sampled, in addition to a suite of boulders from atop the right lateral moraine (Fig. 3; Stop 1.2).

Figure 3. Topographic map of the Angel Lake area (portion of the U.S. Geological Survey 7.5′ Welcome quadrangle), including Stops 1.1, 1.2, and 1.3 (squares). Dashed lines indicate mapped moraine crests, and circles indicate locations of boulders sampled for cosmogenic 10Be surface-exposure dating. Shaded areas indicate the extent of glaciers during the Angel Lake Glaciation.

Pleistocene glacial and pluvial records in northeastern Nevada Before driving to Stop 1.2, note the bouldery surface on the west side of the road just below the break in the distal slope of the terminal moraine. This surface is mapped as Lamoille-age till and is likely part of a moraine that was subsequently buried by till of the Angel Lake Glaciation. This age assignment is based in part on the physical properties of boulders on this surface, which display evidence of intense physical weathering suggestive of prolonged surface exposure. Directions to Stop 1.2 From Angel Creek Campground, retrace the route out of the campground to Angel Lake Road and turn left (Fig. 3). Follow this road through a series of switchbacks up the mountain front and into the cirque holding Angel Lake. Note that this road is extremely steep, narrow, and winding. Use caution when observing roadcuts, and watch for two-way traffic. Park along the road where it straightens out after the final left turn into the cirque (~3.5 mi from turn out of campground). From the road, climb upslope to the south (left) to the crest of the Angel Lake–age right lateral moraine. Stop 1.2. Angel Lake Right Lateral Moraine (41.0246°N, 115.09004°W) The bouldery crest of the right-lateral moraine rises ~35 m above the road and is continuous over a distance of ~0.7 km. It pairs with a left-lateral moraine that is visible to the north across Angel Creek; together, these moraines delimit the vertical extent of the glacier in this valley during the Angel Lake Glaciation (Fig. 3). The frequency of erratic boulders at the moraine crest increases upvalley. Although some of these boulders display clear evidence of substantial surface erosion, many are tall enough and display glacial polish and/or resistant quartz-rich protrusions to render them suitable for cosmogenic 10Be surface-exposure dat-

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ing. We have sampled ten boulders atop this moraine, and processing of these samples is under way. Several additional Pleistocene features are visible from this stop. To the south, a peculiar, bouldery ridge protrudes at a right angle to, and is apparently crosscut by, the right-lateral moraine (Fig. 3). Exposures in this feature along the road indicate that it is composed of till and its surface form suggests that it is a moraine. However, the shape of the glacier that would have constructed this feature is not consistent with the obvious glacier morphology during the Angel Lake Glaciation. Perhaps the Angel Creek glacier flowed southward during a pre–Angel Lake Glaciation to construct this ridge. Alternatively, the ridge may have been constructed by ice in the next valley to the south (South Fork Angel Creek) during a pre–Angel Lake Glaciation. To the north, a well-preserved sequence of recessional moraines is visible as lower ridges inset to the outermost leftlateral moraine along the axis of the Angel Creek valley. These moraines clearly mark the path of ice retreat after the maximum of the Angel Lake Glaciation, with the youngest, innermost moraine impounding the northeast side of Angel Lake. We sampled boulders atop several of these moraines for cosmogenic 10Be surface-exposure dating (Fig. 3). Finally, to the southeast, the Clover and Independence Valleys are visible, situated east of the East Humboldt Range. Together these valleys were occupied by pluvial Lake Clover (Fig. 2), which will be explored on the second day of this field trip. Shoreline ridges of Lake Clover, spanning elevations of 1729–1711 m, are best viewed from this stop during the afternoon hours of sunny days. The southern end of the Clover Valley is now occupied by Snow Water Lake, with a mean elevation of 1707 m. Directions to Stop 1.3 Carefully descend the proximal moraine slope to the road, and drive the remaining stretch up to the final parking lot (straight ahead after the cattle guard). Park the vehicles and walk a short distance to the dam impounding Angel Lake (Fig. 3). Stop 1.3. Angel Lake (41.03380°N, 115.06502°W)

Figure 4. Profile view (to the north) of the Angel Lake and Lamoille terminal moraines from Angel Creek Campground. The Angel Lake moraine is taller, steeper, and more bouldery. The Lamoille moraine is dotted with infrequent, but very large, boulders. The road climbing to Stop 1.2 passes between the two moraines, in front of the bedrock knob visible in the background.

Angel Lake is a natural tarn at 2554 m that was slightly enlarged by a dam built in the early 1900s. The setting is quite scenic, and that quality combined with the easy access from Wells and I-80, makes this a popular spot. Snow melting higher in the cirque during early summer feeds an impressive cataract that enters the lake opposite the dam. A tremendous headwall rising 700 m to Greys Peak encircles the lake on three sides, while a pair of steep-sided lateral moraines extends eastward from the lake along the road. In early summer, the water level is quite high on the dam, but it falls ~2 m as water is released, forming a narrow sandy beach as the summer progresses. From a glacial geology perspective the lake is located at the young end of a well-preserved series of moraines (Fig. 3). The outermost moraines of presumed Lamoille age were viewed at

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Stop 1.1, while the classic Angel Lake moraines were seen from Stops 1.1 and 1.2. Other recessional moraines are clearly visible from the road leading up to the dam, and remnants of inset lateral moraines extend toward the valley axis near the campground as well as on the south side of the road. The youngest end moraine loops down from the north side of the cirque and partially encloses the lake; the artificial dam extends this natural ridge across to the south side to raise the original water level slightly. The position of Angel Lake inside the innermost recessional moraine indicates that the sedimentary record from the lake should contain information about environmental changes occurring in this area since deglaciation. To investigate this record, the lake was cored in June 2007. To select the coring location, a bathymetric map was first created from over 300 depth measurements fixed with a GPS receiver (Fig. 5). The lake has a maximum depth of 11 m, a mean depth of 6.5 m, an area of 53,000 m2, and a volume of 346,000 m3. The watershed area–lake area ratio is 25:1, and assuming a uniform annual precipitation of 0.89 m/ yr (measured at the Hole in Rock SNOTEL site less than 10 km away at a similar elevation on the same side of the range), the flushing rate of the lake is 3.5 times/yr. The bathymetric map (Fig. 5) reveals a deep hole just offshore from the main fluvial input at the western side of the lake. Because the goal was to retrieve a long, undisturbed record,

it was decided to core to the east of the deep hole in slightly shallower water so that episodic flooding events that might deliver large amounts of coarse, reworked sediment to the lake, would not overwhelm the sedimentary record. Care was also taken not to core too close to the dam in case this area was dredged during dam construction. Coring was completed from an anchored platform using a Livingstone corer for the loose surface sediment, and a percussion corer for deeper material. The cores were retrieved in 9.14 m of water, and the composite record extended from the sediment-water interface to a depth of 4.54 m below the lake bottom. The basal sediment of the percussion core contains disseminated shards of tephra, and OSL analysis returned an age (7.94 ± 0.9 ka B.P.) consistent with the Mazama eruption. That constraint, along with 5 accelerator mass spectrometry radiocarbon dates, supports a depth-age model that spans ~7.7 ka B.P. (Fig. 6). Back in the lab, multiple proxies were investigated at 1-cm intervals including: water content, loss-on-ignition (LOI), C:N ratio, biogenic silica content, and grain size distribution (Munroe and Laabs, 2009). LOI and water content show significant transient departures from an overall increasing trend through the record (Fig. 6). Biogenic silica and LOI values are notably above average from 1 to 2 ka B.P. and ca. 7 ka B.P., suggesting a warmer, more productive lake environment. These episodes

Figure 5. Bathymetric map of Angel Lake with depths in meters. Map was created from ~300 GPS-linked depth measurements. The location of the 2007 core is shown by the diamond near the center, just east of the deep hole near the western inlet (arrow).

Pleistocene glacial and pluvial records in northeastern Nevada

Figure 6. Multiproxy time-series from the Angel Lake core. Water content rises through the record reflecting compacting of sediment over time. Loss-on-ignition (LOI, a proxy for organic matter content) also rises, with significant departures from this overall trend. C:N ratio varies reflecting the relative input of terrestrial versus aquatic debris. Biogenic silica values vary widely, perhaps reflecting changes in summer water temperature and/or duration of the ice-free season. Mean grain size (GS) is also highly variable. The analysis described in the text yielded an objective identification of coarse layers inferred to record wet avalanches onto the lake ice. The frequency of these events was highest in the middle Holocene, and ca. 3500 and 1000 yr B.P. The thicker coarse layers are visible as light gray bands on the X-radiograph. The depth-age model was based on five accelerator mass spectrometry 14C dates, the year A.D. 2007 for the core surface, the presence of tephra assumed to be from the Mazama eruption near the base, and an optically stimulated luminescence (OSL) analysis on the basal sediment.

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of increased productivity are broadly consistent with times of increased chironomid-inferred July air temperatures at Stella Lake in Great Basin National Park, 250 km to the south (Reinemann et al., 2009). In particular, the interval of peak warmth at Stella Lake ca. 5.5 ka B.P., is synchronous with an interval of particularly high LOI values (Fig. 7). Higher LOI values in the later Holocene overlap with times of reconstructed high water levels in Ruby Lake to the southeast, as well as Pyramid Lake and Lake Tahoe at the western end of the Great Basin (Fig. 7). Together these similarities suggest a general increase of both relative moisture and productivity in high-elevation lakes in the late Holocene. Mean grain size in the Angel Lake record is highly variable, with spikes to locally high values reflecting delivery of clastic debris to the coring site by high-energy events (Fig. 6). One possible interpretation is that these coarse layers represent floods on the inlet stream draining into the lake. However, because the core was retrieved from the opposite side of the depocenter from the inlet in water ~2 m shallower than the deepest part of the basin, it is unlikely that these clastic layers are evidence of direct fluvial inputs. Instead, these layers are interpreted to represent depositional events during the winter and spring when run-out from snow avalanches and slushflows could travel across the ice to the coring site (Munroe and Laabs, 2009). To objectively identify these discrete events in the record, a Gaussian smoothing function was run through the grain size time series and avalanche events were defined as peaks rising above the local background level. Results from this analysis indicate that the frequency of avalanches was below average from 1.8 to 3.2 ka B.P., which

overlaps the extended low in LOI. In contrast, avalanches were quite common, up to 2 times the long-term average, from 0.5 to 1.5, 3.2 to 3.8, and from 4.5 to 6.5 ka B.P. The oldest interval is synchronous with a prolong interval of warm temperatures at Stella Lake, as well as lowstands in Ruby Lake, Pyramid Lake, and Lake Tahoe (Fig. 7). Together these fluctuations may reflect changes in winter/spring moisture delivery to the northeastern Great Basin. Alternatively, increases in avalanche frequency may indicate more common rain-on-snow events or episodes of rapid warming during the spring thaw, both of which could stimulate wet avalanches and slushflows. Directions to Elko Return to I-80 via the Angel Lake Road (NV-231). Use caution on the steep descent, particularly if driving a fully loaded van. Use of lower gear is recommended in order to prevent brakes from overheating. Pass under the highway and turn left to access I-80 westbound. Travel 47 mi to Elko and take exit 303. Turn left under the highway, then right on Idaho Street. Turn right into the Red Lion Hotel. DAY 2. ELKO TO ELKO: PLUVIAL LAKES FRANKLIN AND CLOVER (Total driving distance: ≈217 mi/350 km over Harrison Pass, 247 mi over Secret Pass.) Day 2 Overview Day 2 focuses on the record of pluvial Lakes Franklin and Clover on the east side of the Ruby–East Humboldt Mountains (Fig. 2). We will visit several sites illustrating aspects of the last highstand and regression of both lakes, and will receive an overview of the Ruby Lakes National Wildlife Refuge from Refuge staff (Fig. 8).

Figure 7. Comparison of Angel Lake record with other records from the region. A—Stella Lake in Great Basin National Park (Reinemann et al., 2009); B—Ruby Lake (Thompson, 1992); C—Lake Tahoe (Benson et al., 2002; Lindstrom, 1990); D—Pyramid Lake (Benson et al., 2002); E—Blue Lake (Louderback and Rhode, 2009). Intervals of high loss-on-ignition (LOI) at Angel Lake generally line up with times of warm temperatures at Stella Lake in Great Basin National Park. High LOI also corresponds to a greater frequency of clastic layers. Intervals of fine grain size (GS) are out of phase with high LOI. See text for discussion.

Directions to Stop 2.1 From the Elko Red Lion hotel, there are two options for reaching the Ruby Valley. The first is to head south along the west side of the Ruby Mountains before crossing into the Franklin Valley at Harrison Pass. This route is shorter, but the Harrison Pass Road is steep and winding, and parts of it are unpaved, so it may not be possible to cross there in late May. The alternative route is 30 mi longer, but ascends gentler grades to cross the range at Secret Pass, and remains on paved roads. For the purpose of flexibility, and to aid future users of this guidebook, both sets of directions to Stop 2.1 are presented here. Harrison Pass Route to Ruby Lakes National Wildlife Refuge (62 mi): Turn right on Idaho Street from the Red Lion Hotel. Turn left on 5th Avenue and follow south across the Humboldt River. Continue to follow 5th Avenue as it curves left and becomes NV-227 toward Spring Creek. After ~7 mi, turn right at the light on NV-228. Follow this road south for 30 mi along the western side of the Ruby Mountains through Jiggs and then veer left toward Harrison Pass. Climb toward the pass on U.S. Forest

Pleistocene glacial and pluvial records in northeastern Nevada

Figure 8. Map of the Lake Franklin region with outlines of the lake at five different elevations (in meters). Field trip stops are identified, and the route of the trip (from Fig. 2) is shown.

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Service (USFS) Road 113. Note that the last few switchbacks on this road are unpaved, and the road is not maintained in the winter. A pull-out on the right in Harrison Pass provides a nice view of a landscape featuring eroding outcrops of the Oligocene Harrison Pass Pluton (Howard, 2000). To the east, the Franklin Valley is visible. Descend the east slope of the Ruby Mountains in a narrow canyon and reach Ruby Valley Road at the base of the mountains. Turn right for the Ruby Lakes National Wildlife Refuge (Fig. 8). Secret Pass Route to Ruby Lakes National Wildlife Refuge (92 mi): Turn left on Idaho Street from the Red Lion Hotel, make the first left on East Jennings Way, and then a quick right to enter I-80 eastbound. Travel 17.6 mi east to exit 321 and follow NV-229 toward Ruby Valley. This road leads over Secret Pass between the northern end of the Ruby Mountains and the southern end of the East Humboldt Range. After 35.6 mi from I-80, continue straight on NV-767 (Ruby Valley Road) along the base of the eastern slope of the Ruby Mountains. This road will meet the route coming over Harrison Pass, and continuing straight will lead to the Ruby Lake National Wildlife Refuge (Fig. 8). The headquarters of the Refuge is on the right ~9 mi south of the intersection of the two routes along a well-maintained dirt road. Stop 2.1. Ruby Lake National Wildlife Refuge and Pluvial Lake Franklin (40.22654°N, 115.48932°W) The Ruby Lakes National Wildlife Refuge (RLNWR) was established in 1938 because of its critical location at the crossroads of several major corridors for migrating birds. Today the refuge covers ~40,000 acres, almost half of which is wetlands of the Ruby Marshes. More than two hundred natural springs at the base of the Ruby Mountains deliver water to the marshes, and the Refuge staff manages water levels within different sections of the marsh with a system of dikes, pumps, and canals. A self-guided auto-tour passes through the marshes along several of these dikes, providing abundant opportunities for birdwatching. Maps of the auto tour route are available at the Refuge headquarters. The nearby Gallagher State Fish Hatchery also takes advantage of the water from the springs. The hatchery, which was built in 1940 and substantially updated in the late 1960s, primarily raises trout that are stocked in lakes and streams around northern Nevada. The RLNWR occupies a valley that formerly held pluvial Lake Franklin, one of the largest pluvial lakes in the Great Basin after Bonneville and Lahontan (Mifflin and Wheat, 1979). At its highest level (~1853 m) Lake Franklin inundated the RLNWR with over 35 m of water, covered more than 1000 km2, and extended northward through the Franklin Valley, eastward across Dry Lake Flat, and into the North Butte Valley (Fig. 8). All that is left of Lake Franklin today is the shallow Ruby Marshes and an ephemeral lake on the floor of Franklin Valley. As is the case with most pluvial lakes other than Bonneville and Lahontan, the record of Lake Franklin has received only limited attention. Shoreline features recording a highstand of Lake Franklin were noted by the earliest geological explorations in the area, but were inaccurately considered evidence of a much

larger lake (Simpson, 1876). Later expeditions correctly realized that these features reflected a lake in the Franklin Valley (Gilbert, 1890; Russell, 1885). Several decades later, the name Lake Franklin was assigned to this pluvial lake, taken from the modern ephemeral lake on the floor of the Franklin Valley (Sharp, 1938). Despite this early recognition, only two studies have focused specifically on the history of Lake Franklin and attempted to develop chronologies of its former fluctuations. The first used lacustrine sediment cores from the vicinity of the Ruby Marshes to reconstruct the past 40,000 years of water level changes (Thompson, 1992). Thompson (1992) studied two cores: one drilled to a depth of >7 m at the edge of the modern marsh, and another collected in ~2 m of water just offshore from the first coring site. By comparing the relative abundance of different aquatic palynomorphs, Thompson was able to reconstruct how salinity of the marshes had changed over time, which was considered a signal of water depth. The overall conclusion is that relative fresh and deep water covered the coring site between ca. 18,500 and 15,400 14C yr B.P. These dates calibrate to ca. 24,500–19,500 and 20,100–16,700 calendar yr B.P., suggesting that the highstand of Lake Franklin was synchronous with MIS-2 and occurred during the final rise to the Bonneville highstand in the Bonneville Basin (Oviatt et al., 1992). A more complete investigation of the history of Lake Franklin was completed as a Ph.D. dissertation at the University of Utah (Lillquist, 1994). Lillquist (1994) undertook extensive mapping of shoreline features within the area flooded by Lake Franklin, and was able to assign ages to many of these shorelines through radiocarbon dating of shell fragments retrieved from natural exposures and artificial excavations in beach ridges and adjacent lagoons. Lillquist (1994) also recognized several significant shorelines created by Lake Franklin at elevations below the highstand, and constructed a hydrograph charting several cycles of lake regression and transgression during the overall desiccation of the lake from MIS-2 into the Holocene (Fig. 9). The highstand of Lake Franklin at 1853 m is constrained in Lillquist’s hydrograph by a pair of dates on shells retrieved from auger holes drilled in lagoons. These dates indicate that Lake Franklin built its highest shoreline between 16,800 ± 130 and 15,070 ± 100 14C yr B.P. (20,000– 18,300 yr B.P.; note: for clarity all calibrated dates are reported as the midpoint of the most probably 2-σ calibration range). A date of 15,020 ± 240 14C yr B.P. from a small quarry directly in a beach ridge indicates that water level had fallen to 1843 m by 18,200 B.P. A significant regression to ~1823 m occurred by 14,650 ± 340 14C yr B.P. (17,800 B.P.) based on evidence exposed at the Franklin River Bridge (Stop 2.3). Lake level rose again after this time reaching 1836 m and then 1840 m, with a brief regression to 1826 m in between. The final regression of the lake began at the 1840 m shoreline ca. 12,720 ± 110 14C yr B.P. (15,000 B.P.), and the last pair of shorelines was constructed at ~1820 m by 11,500 14C yr B.P. (13,000 calibrated B.P.) (Lillquist, 1994). The work of Lillquist (1994) is significant for the perspective it provides on lake level dynamics during the last GIT: clearly, Lake Franklin didn’t regress monotonically from its 1853 m

Pleistocene glacial and pluvial records in northeastern Nevada

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Figure 9. Radiocarbon dates for shorelines of Lake Franklin. The hydrograph of Lillquist (1994) is shown, along with a revised composite version taking into account newly acquired radiocarbon dates. Dates directly from beach ridges are shown as diamonds, while (minimum-limiting) dates from lagoons are shown as squares. Open symbols represent dates on reworked material not used in constructing the hydrographs. Elevations of the highstand, the 1843 m shoreline, the Ruby Lake National Wildlife Refuge (RLNWR) quarry, and the Franklin River Bridge site are noted.

highstand. But as noted earlier, the import of these lake level oscillations during overall regression is unclear. What magnitude of climate changes do they represent? Do they reflect the effects in the Franklin Valley of a regional climate forcing, or were they driven by something more local and basin-specific? Did these changes occur synchronously with climate shifts during the last GIT elsewhere in the Great Basin? Ongoing research by the authors is intended to address these questions by further developing the chronology of Lake Franklin fluctuations, and comparing it with records from the unstudied pluvial Lake Clover, as well as glacial moraines deposited during the Angel Lake Glaciation and subsequent deglaciation.

northeast from the mountain front and is bordered to the southwest by a large alluvial fan draining off the southern Ruby Mountains. Lillquist (1994) inferred that the fan served as a major source of sediment for the spit, and concluded that northward sediment transport was dominant, with a minor component of transport to the south and east. The north wall of the quarry exposes ~4 m of well-rounded, crudely stratified gravel and coarse sand (Fig. 10). The sediment

Directions to Stop 2.2 Continue south from the Refuge headquarters past the Gallagher State Fish Hatchery and turn left on (gravel) Brown Dike road (Fig. 8). Park on the right and walk a short distance down into the RLNWR gravel pit for Stop 2.2. Stop 2.2. Ruby Lake National Wildlife Refuge Gravel Pit (40.17722°N, 115.48968°W) Please note that the pit is not open to the public. Visitors wishing to view the sediment exposed at this location should check with RLNWR staff at the Refuge headquarters before entering the pit. The excavation at this stop was a source of fill used in creating roads and dikes within the RLNWR. The pit is located within a compound spit formed by transport of sediment by waves and currents when Lake Franklin stood at the 1830 m level, ~23 m below the highstand elevation (Fig. 8). The spit projects east-

Figure 10. Exposure in the north wall of the Ruby Lake National Wildlife Refuge gravel quarry in June 2010. The face stands ~4 m high and exposes cross-bedded, indurated, coarse gravelly sand interpreted as a beach environment. Shells dated from this section constrain the elevation of Lake Franklin to 1830 m between 20,000 and 19,000 yr B.P.

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is moderately well indurated and is locally abundantly fossiliferous. Lillquist (1994) interpreted these sediments as a series of lagoon and foreshore/beach units, with the lagoons represented by marly silty sands and the beaches by marly sandy gravels. Changes in the exposure over time due to continued excavation make it impossible to determine exactly how the modern exposure relates to that described by Lillquist (1994). However, the exposure in 2010 was dominated by a coarser facies of indurated gravelly sand, suggesting that it was located solidly in the former beach environment, rather than in a lagoon. Lillquist (1994) reports three previously unpublished radiocarbon ages from this site collected by D.R. Currey and B.G. Bills. These ages on shells, tufa, and marl from lagoonal environments within this spit complex, range from 12,030 ± 140 14C yr B.P. (13,960 B.P., shells) to 18,030 ± 1050 14C yr B.P. (21,460 B.P., marl). In addition to these, Lillquist (1994) reported ages of 14,360 ± 150 14C yr B.P. (17,500 B.P.) and 12,870 ± 140 14C yr B.P. (15,650 B.P.) on shells from lagoonal facies at this site. In 2010, the authors obtained two more accelerator mass spectrometry radiocarbon analyses on gastropods collected directly from upper and lower sections of the beach gravels exposed in the north wall of the pit. The ages on these samples are 15,750 ± 140 and 16,750 ± 70 14C B.P. These ages, which are in stratigraphic order, calibrate to ca. 19 and 20 ka B.P., respectively. Dates on shell material from lagoons should be viewed as providing a minimum limiting age on formation of the neighboring ridge because, as Lillquist (1994) notes, the lagoon could remain a locally moist environment capable of harboring mollusks long after the water level has dropped from the level of the ridge. In contrast, dates on shells recovered directly from a beach ridge provide a more direct constraint on the age of the highstand responsible for ridge formation. Viewed in this way, the two new dates for the RLNWR indicate that the water level in Lake Franklin stood at ~1830 m ca. 20–19 ka B.P. The younger dates obtained from this section by Currey, Bills, and Lillquist were all from lagoonal facies, which could have held water at later times when the main lake rose and fell again below the 1830 m level. The two new dates from the RLNWR quarry require a change in the Lake Franklin hydrograph developed by Lillquist (1994). At first glance, the results from the RLNWR quarry and Lillquist’s two sites constraining the Lake Franklin highstand appear incompatible because they suggest the water level was at 1853 m and 1830 m simultaneously. On closer inspection, however, an explanation is revealed by the fact that Lillquist’s dates for the highstand are from lagoons, which as noted earlier should be viewed as minimum limiting ages. Combining the dates from lagoons behind the highstand beach ridge and the new shell dates directly from the ridge at the RLNWR quarry, it appears that the highstand of Lake Franklin occurred earlier than Lillquist (1994) suggested, perhaps closer to 21 ka B.P. before the water level fell to the 1830 m level by 20–19 ka B.P. (Fig. 9). This interpretation places the Lake Franklin highstand solidly in the middle of MIS-2, and also aligns cleanly with the duration of the highstand inferred by Thompson (1992). Interestingly, a drop (~40 m) in

the elevation of Lake Bonneville known as the Keg Mountain Oscillation has been reported from approximately the same time as the drop in Lake Franklin to the 1830 m level (Oviatt et al., 1992; Burr and Currey, 1988; Currey and Oviatt, 1985). Directions to Stop 2.3 Turn vehicles around, return to Ruby Valley Road, and head back to the north (right) to exit the Refuge. A quarter-mile past the intersection with the Harrison Pass Road coming in from the left, turn right on the CCC Road (Fig. 8). This dirt road leads along the crest of a narrow ridge that separated Lakes Franklin and Ruby when the water level fell below ~1830 m. Lillquist (1994) inferred that this segmentation feature formed through simultaneous eastward progradation of a spit from the western shore of the lake, and westward progradation of the alluvial fan emanating from Ruby Wash on the eastern shore. This interpretation is based on the overall geomorphology of the feature, as well as the dominance of granitic lithologies derived from the Ruby Mountains in the western half of the ridge, and carbonates from the Maverick Springs and Medicine Ranges in the eastern half. The main road, which is easy to follow, leads across to the east side of the valley, then turns northward and follows the 1830 m shoreline around the distal end of the large Ruby Wash fan. The road then rises up to the higher shorelines at 1843 and 1850 m and curves back around to the north. Continuing around the west side of the Dry Lake Flat, the CCC Road follows a substantial compound ridge with a crest elevation between 1843 and 1846 m. Lillquist (1994) notes the presence of superimposed shoreline features on this ridge, suggesting that the water level reached this elevation on more than one occasion. The road once again curves to the north near Murphy Well where Lillquist (1994) obtained his highest radiocarbon date directly from a beach ridge (18,200 yr B.P. at 1843 m). Watch for a sign marking the road to Franklin River Bridge on the left ~26 mi after turning onto the CCC Road. Take this turn and follow 6 mi to the site of the former bridge (now impassable). Park along the road on the east side of the bridge for Stop 2.3. Stop 2.3. Franklin River Bridge (40.52899°N, 115.20827°W) At this stop, we will review geomorphic and stratigraphic evidence for a major regression in Lake Franklin following the formation of the 1843 m shoreline. The road leads from this shoreline to the Franklin Bridge site across the eastern half of a broad ridge that cuts off the northern part of the Lake Franklin Basin. Lillquist (1994) interpreted this ridge as a compound spit/bayhead bar that has been breached by the combined flow of Withington Creek and the Franklin River. Lillquist (1994) described an exposure of the stratigraphy at this site as ~100 cm of gravelly sand unconformably overlying 50 cm of fine clay, over 50 cm of clayey sand. The fine clay, which contains ostracodes and mollusk fragments, was interpreted as a deepwater deposit, and the overlying gravelly sand as a beach. A date of 14,650 ± 340 14C yr B.P. (17,800 B.P.)

Pleistocene glacial and pluvial records in northeastern Nevada from the base of the gravelly sand provides a constraint on the timing of the lowstand responsible for the beach. Lillquist (1994) reported two other dates from this section: 7320 ± 90 14C yr B.P. (8150 B.P.) from carbonate concretions in the lowest clayey sand, and 16,880 ± 510 14C yr B.P. (20,156 B.P.) from an extended counting analysis on small shell fragments from near the top of the section. The lower date is clearly too young and likely represents the time of formation of the concretions, rather than deposition of the sediment. The upper date is out of stratigraphic order, and given the fragmented nature of the shells, likely represents reworking of older shells from a higher shoreline down to the Franklin Bridge site by the Franklin River at the time of the lowstand. In 2010, the exposure revealed ~150 cm of coarse, gravelly sand unconformably overlying a sticky fine silty loam with 13% clay (Fig. 11). The authors obtained a date of 18,300 ± 95 14C yr B.P. (21,850 B.P.) on shells from the gravelly sand, 145 cm below

Figure 11. Stratigraphic section exposed at the east end of an excavation east of the Franklin River Bridge in June 2010. The section exposes ~100 cm of bedded, fossiliferous, coarse sandy gravel interpreted as a beach deposit. This unit sits unconformably on a silty-clay layer (not shown) interpreted as a deep-water facies. Shell samples were collected from the beach gravel (near tip of shovel handle) and a shielded sample for optically stimulated luminescence analysis was collected from the same level (excavation at left).

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the modern surface. We interpret this as additional evidence of reworking of older material downstream to the location of this beach. As a cross-check on this interpretation, we also collected a sample for OSL dating from the same layer where we dated shell fragments. Results from this analysis should provide some clarification of the mixed ages from this site. The importance of this regression to the Franklin River Bridge site is unclear; however, it is obvious that the climate changed significantly to reduce the area and volume of the lake so dramatically. Calculations in a geographic information system (GIS) reveal that Lake Franklin lost 53% of its area (1000– 467 km2) and 89% of its volume (17–2 km3) during the drop from the 1843 shoreline (ca. 18,200 B.P.) to the Franklin Bridge (ca. 17,800 B.P.). Following this lowstand, the lake elevation rose again, returning to the 1843 m shoreline by ca. 16,800 B.P., given a new date of 13,400 ± 75 14C yr B.P. (16,400 B.P.) obtained by the authors on a beach ridge in Ruby Wash. This submergence of the bayhead barrier/spit at Franklin River Bridge was likely responsible for the muted form and broad appearance of the feature today. Auger excavations west of the bridge and south of the road reveal ~20 cm of fine silt overlying coarse, rounded gravel which may represent deeper water sediments that accumulated over the beach facies during this transgression. Reoccupation of the 1843 m shoreline is consistent with the presence of superimposed beach features at the 1843 m level noted by Lillquist (1994). The 1843 m shoreline marks the general elevation of the division between the main Franklin Valley and the North Butte– Dry Lake Flat sections of the basin. Lillquist (1994) suggested that spillover of water from the Franklin Valley into these basins to the east could have helped maintain the lake at this elevation. Calculations in a GIS support this interpretation. The area of Lake Franklin at the 1843 m elevation was 1000 km2, but 140 km2 of that was contained within the shallow arm extending east into the North Butte Valley (Fig. 8). Thus, the rise of the lake from just below, to just above, the 1843 m level involved an instantaneous increase in surface area of 16%. Such a substantial increase in area available for evaporation could have exerted a strong limiting influence on continued water level rise. Directions to Stop 2.4 Retrace the route back to the west and up to the CCC Road, note mileage, and turn left to continue north (Fig. 8). The road generally follows just below the shoreline at 1850 m, and in several places provides great overviews of the 1843 m shoreline to the west (Fig. 12). After passing Hequy Well (~5 mi), where Lillquist (1994) obtained a date of 15,070 ± 100 14C yr B.P. from a lagoon east of the road, the road passes near a small cuspate spit formed as a paired of tombolos extended out to a small bedrock island while the lake stood at the 1840 m level (Fig. 13). North of this point, the shoreline transitions from an isolated ridge to one plastered on the side of a bedrock upland. After ~8 mi, park alongside the road and hike downslope to the west to reach the bedrock outcrop for Stop 2.4.

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Stop 2.4. Wave-Cut Platform at Lake Franklin Highstand (40.60479°N, 115.13465°W)

Stop 2.5. Deep Gravel Quarry on CCC Road (40.64649°N, 115.13791°W)

This stop provides an expansive overview across the northern end of the Franklin Valley, as well as access to a local erosional feature produced when the lake was at the 1850 m highstand (Fig. 13). At this location, Lake Franklin was nearly 12 km wide when the highstand shoreline was occupied, and maximum fetch for southerly winds exceeded 50 km. Production of substantial waves was possible in a lake of this size, and the flat surface eroded into the limestone bedrock at this location appears to be a wave-cut platform. Most of the highstand shoreline segments of Lake Franklin mapped by Lillquist (1994) are depositional, but 40% were classified as erosional, and Sharp (1938) also commented on wave-cut features in the basin. The surface of this outcrop is heavily weathered, but the overall effect of wave planation is visible. The elevation here is 1848 m, so water depths were on the order of 2–5 m during the highstand. Reconnaissance along the bedrock outcrop to the east of the road reveal rounded cobbles in the apron of sediment at the base of the bluff, suggesting that waves may have cut the bluff as a sea cliff. Lillquist (1994), however, notes that the presence of numerous faults in this area complicate identification of linear features related directly to lake activity.

This final stop in the Franklin Valley focuses on a gravel pit where material was removed from a compound beach ridge for road construction. Lillquist (1994) describes the stratigraphy in a former exposure within this excavation where interbedded sandy gravel and gravelly sand were visible. Beach deposits of the highstand shoreline were originally present just east of this site, but have been largely mined away or disturbed by equipment. This sandy gravel was underlain by a gravelly sand interpreted as nearshore deposit representing regression from the 1850 m highstand. A thin layer (~10 cm) of sandy clay beneath the gravelly sand was interpreted as offshore clay representing the (short?) time when the water level was above this site, forming the highstand beach. This unit is underlain by another thin layer of gravelly sand representing foreshore deposits of the transgression to the highstand level. The lowest unit of the exposure was sandy gravel heavily cemented by carbonate, and Lillquist (1994) reported the carbonate morphology to be equivalent to Stage IV (Birkeland et al., 1991). Given this level of development, he concluded that this gravel represents a much older beach deposited when an ancestor of Lake Franklin rose to this level earlier in the Pleistocene. Lillquist (1994) did not date this layer, but suggested that its age could approach 750,000 yr given the similarity of carbonate development to a site in southern Utah of that age.

Directions to Stop 2.5 Hike back up to the vehicles and continue northward along the CCC Road. After 3 mi, the road passes through an old gravel quarry. Park near the north end of the quarry for Stop 2.5.

Figure 12. The 1843 m shoreline below the CCC Road in northeastern Franklin Valley (view to the north). The broad crest of this shoreline extends northward away from the viewer, and is visible curving off to the northwest. This ridge is paralleled at a lower elevation to the left (west) by another shoreline representing the 1840 m water level. The crests of both ridges are accentuated by lighter-colored vegetation. See Figure 8.

Figure 13. Aerial photograph overlain on topographic map showing part of the northeastern Franklin Valley (portion of the U.S. Geological Survey 7.5′ Smith Well quadrangle). The CCC Road followed by the field trip route is visible as the lightcolored line. The wave-cut platform at Stop 2.4 is highlighted. Note the location of the platform where the highstand shoreline comes in contact with a bedrock-cored upland. The small cuspate spit connected to a bedrock outcrop at the 1840 m shoreline is also identified, along with numerous other shorelines (from 1840 m to 1830 m) visible at the southern extent of the map.

Pleistocene glacial and pluvial records in northeastern Nevada In 2010, the authors obtained a shell sample from this cemented layer that was suitable for radiocarbon dating. The shell returned a date of 42,300 ± 480 14C yr B.P. (45,600 B.P.), indicating that the beach certainly predates the MIS-2 highstand, even if Lillquist’s estimate is too old. It is unclear how this date should be interpreted. Taken at face value, it suggests that a lake stood near the level of the MIS-2 highstand during MIS-3. Little is known about pre–MIS-2 lakes in the Great Basin, although older shorelines have been reported from some basins. In most cases, the ages of these features are unknown; however, work with a long sediment core from the Bonneville Basin suggests, not surprisingly, that deep lakes formed during glacial periods (Oviatt et al., 1999). Extrapolating this conclusion westward, it seems unlikely that Lake Franklin would have stood near its highest level for MIS-2 during the interstadial of MIS-3. Maybe this sample failed to yield an accurate age because of aragonite to calcite recrystallization. In this scenario, the age should be viewed as a minimum limit, and the actual age of the carbonate-cemented gravel would be older than 45,600 years, perhaps representing deposition by a high lake during MIS-4 or even MIS-6. Directions to Stop 2.6 Continue north on the CCC Road a short distance to intersect NV-229. Turn right on this paved road and climb over the divide out of the Franklin Valley (Fig. 8). At the intersection with U.S.-93, set trip odometer to zero, turn left (north), and descend into the Clover Valley (Fig. 14). The highstand of pluvial Lake Clover at 1729 m is crossed ~5.3 mi north of this intersection, although no shoreline feature is visible. The road continues down into the basin and passes near the western end of the Snow Water Lake playa (Bureau of Land Management [BLM] sign). An impressive set of dunes rises ~10 m above the playa surface to the southeast, while other dune ridges are visible along the north side of the playa in the distance. Continuing north on U.S.-93 watch for a right turn entering a gravel pit marked by a small BLM sign for “Tobar” ~19 mi north of the U.S.-93 intersection. Turn here, pass through the gravel pit, and continue to follow the dirt road as it veers south following the railroad tracks into the Clover Valley. At a prominent intersection (25.5 mi) stay to the right on the west side of the tracks. The road continues southward, drops below the 1729 m shoreline, and runs parallel to a prominent ridge at 1716 m to the left. This ridge extends southward as a tombolo toward an upland known as Black Ridge. Follow the road to the crest of this upland (~29 mi) and park along the left after the fence marking the end of private property around a house trailer for Stop 2.6. Stop 2.6. Black Ridge (40.85248°N, 114.87100°W) Several shorelines of Lake Clover are preserved in the vicinity of Black Ridge, a low-relief, bedrock- (or gravel-) cored ridge that was an island near the north-central shoreline during the highstand of Lake Clover (Fig. 14). The highest elevation of Lake

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Clover (1729 m) is represented by a nearly continuous shoreline ridge along the broad alluvial fan in the north-central sector of the basin. Shoreline ridges representing regressive phases of the lake are also preserved here; the most continuous are at elevations of 1725, 1722, and 1716 m. Although the ages of the Lake Clover shorelines are not known, we have collected fossil mollusk shells from multiple sites in the valley that will provide 14C age limits on shorelines at 1729, 1725, and 1716 m. Directions to Stop 2.7 From the summit of Black Ridge continue south on the gravel road, dropping immediately below the shoreline from the Lake Clover highstand (Fig. 14). Depending on the light conditions, several subtle shoreline benches are visible as the road drops down to cross the (likely dry) channel known as “The Slough” that leads to a terminal basin in the Independence Valley east of the road. The road becomes finer and locally rutted at these lowest elevations, but should be passable unless it has recently rained. Approaching the northern extent of Spruce Mountain Ridge, the road rises up onto an impressive complex of shoreline features forming a series of nested cuspate spits. Stay left at the BLM signboard (33 mi), and then make a very sharp right on a two-track heading back to the west (34 mi). Park alongside the road for Stop 2.7. Stop 2.7. Ventosa (40.79518°N, 114.84453°W) Shorelines at the south-central end of Lake Clover are visible at this stop. Here, the most prominent shoreline feature (quite obvious on a topographic map or aerial photograph) is a V-shaped cuspate spit at 1715 m (Fig. 15). The gravel road trending northwest-southeast is built upon the eastern limb of this feature, while aerial photography reveals multiple ridges along the western limb of the spit that are cut by the northeast-southwest–trending road. The center of the triangle was a lagoon when Lake Clover stood at the 1715 m level. To collect samples suitable for OSL dating from this landform, and to investigate its internal stratigraphy, a trench was excavated across the northwest-southeast–trending portion of the spit in 2008 with a backhoe. The trench exposed the upper 2 m of sediment over a distance of ~20 m along the back slope of the ridge, perpendicular to the trend of the spit (Fig. 16). The trench revealed that the spit is composed of thin beds of open-work gravel, gravelly sand and coarse sand, with clast sizes ranging from pebbles to small cobbles. Most beds are continuous over less than 5 m, and gently dip (<5°) toward the depression west of the spit. Clasts within 1 m of the surface display thin (<2 mm thick) secondary carbonate coatings on their undersides (Fig. 16). No shells were found in the sediments exposed in the trench; however, OSL ages for the trench sediments range from 11.35 ± 2.76 to 8.68 ± 2.07 ka B.P. (Laabs and Munroe, 2009). We view these as minimum age limits because they are inconsistent with hydrographs for lakes in the nearby Franklin and Bonneville basins (e.g., Lillquist, 1994; Oviatt, 1997). Lake Bonneville fell

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below its outlet threshold by ca. 17,300 B.P. and experienced an overall decline in shoreline elevation after ca. 15,000 B.P. Based on the updated chronology of Lake Franklin shorelines presented earlier, water levels were falling after this time there too. Additionally, uncertainties in the time-integrated in situ water content of the shoreline deposits may significantly affect OSL ages. Although the stratigraphy of the spit sediments suggests continuous deposition during the lake stand at 1715 m, it is unclear how long the water level of Lake Clover remained at the elevation of

the spit after it was deposited, or if the water level rose back up to this level at some point after the spit was first formed. Either way, the modern in situ water content could underestimate the longterm average, which would result in OSL ages younger than the true age of the spit. We are testing for systematic errors in these OSL ages by applying 14C dating to the same shoreline ridges. The orientation of the V-shaped cuspate spit and other shoreline features at this elevation suggests a prevailing southerly wind direction during the regressive phases of Lake Clover. Waves

Figure 14. Map of the Lake Clover region with outlines of the lake at five different elevations (in meters). Field trip stops are identified, and the route of the trip (from Fig. 2) is shown.

Pleistocene glacial and pluvial records in northeastern Nevada approaching from the south, southwest, or southeast were necessary for formation of the cuspate spit at this site (Fig. 15), the tombolo near Black Ridge (Fig. 14), and the long spit at the northeast end of the lake basin (Fig. 14), all of which are at ~1716 m. Although the age of this shoreline is not known, a southerly prevailing wind direction is contrary to westerly wind fields inferred from spit orientations and lake-circulation modeling for the regressive phases of Lake Bonneville, particularly for the Provo shoreline that spans 17.3–14.5 ka B.P. (Oviatt, 1997; Jewell, 2010). However, southerly winds are prevalent in the modern wind field of the northeastern Great Basin, becoming most powerful during storm events associated with low-pressure systems. If the modern wind field was established during the regressive phases of Lake Clover, then it is likely that the ~1716 m shoreline was formed at a time when regional circulation was no longer strongly affected by North American ice sheets, sometime after the global LGM, 20–19 ka. Higher shorelines of Lake Clover are visible at this site too, as nearly continuous, low-relief ridges on the alluvial fan to the south (Fig. 15). The highest shoreline at 1729 m is composed of gravels and sands similar to those comprising the V-shaped cuspate spit. A 1-m-deep trench was excavated in this shoreline and samples were collected for OSL dating in 2009. Once again, OSL ages for this shoreline (ranging from ca. 4 to 6 ka) are likely too young. As a possible remedy, an extensive augering campaign on beach ridges and lagoons throughout the Clover Valley in 2010 yielded several samples of mollusk shells suitable for radiocar-

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bon dating, including several from the 1729 m highstand. Results from these analyses will hopefully help refine the ages of the main Lake Clover shorelines. Directions to Elko Retrace the route back out to U.S.-93 and turn right to head north to Wells. Pass under the highway and turn left to enter I-80 westbound. Continue west to Elko (47 mi) and take exit 303 to return to the Red Lion Hotel. DAY 3. ELKO TO LOGAN: THE LAMOILLE TYPE LOCALITY AND LAMOILLE CANYON (Driving distance: ≈260 mi/420 km.) The final day of this trip will focus on the glacial geology of the Ruby Mountains (Fig. 2). After checking out of the Red Lion Hotel in the morning, we will head directly to the western side of the Ruby Mountains to visit the type locality for the Lamoille Glaciation, and the spectacular Lamoille Canyon. After lunch we’ll head north to I-80 and then retrace our route to Logan across northwestern Utah. Please note that the owners of the Ruby Dome Ranch have kindly provided access to their land for this field trip to visit the Lamoille type locality and Seitz Canyon. This permission is not extended to future followers of this field trip route unless specific arrangements are made with the ranch owners. Thus, Stops 3.1 and 3.2 should not be attempted by anyone seeking to replicate

Figure 15. Shorelines of pluvial Lake Clover at the Ventosa area viewed at Stop 2.7 (portion of the U.S. Geological Survey 7.5′ Ventosa quadrangle). The highstand shoreline (1725 m) hugs the northern extent of the bajada extending from Spruce Mountain Ridge (south of figure). The next lower shoreline (1719 m) exhibits a transition to northward progradation. The lowest shoreline, at ~1715 m, forms a prominent V-shaped cuspate spit indicating that prevailing wind directions were driving sediment transport northward when the lake stood at this level. Circles indicate sites where trenches and/or auger holes were dug to sample deposits for optically stimulated luminescence and/or 14C dating.

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this exact itinerary. Fortunately, extensive views of the Lamoille deposits are visible from the road heading toward Lamoille Canyon that bisects the moraine belt. Directions to Stop 3.1 From the Red Lion Hotel, turn right on Idaho Street and then left on 5th Avenue. Follow 5th Avenue as it curves left to become NV-227 toward Spring Creek. This road continues for ~19 mi toward the Ruby Mountains, providing tremendous views of the steep glacial valleys south of Lamoille Canyon. Depending on the light conditions, moraines of the Angel Lake Glaciation (arcuate ridges crossing the valleys at higher elevations) and Lamoille Glaciation (continuous lateral moraines at lower elevations) may be visible. After 19 mi, turn right on Lamoille Canyon Road and head south toward the range front. The road crosses a compound outwash surface deposited during multiple Pleistocene glaciations, and reaches deposits of the Lamoille type locality near the mouth of Lamoille Canyon (Fig. 17). The Ruby Dome Ranch is accessed through a gate on the right just inside the moraine belt. Close the gate behind the vehicles and descend the slope to the left. Park along the road at the base of descent, and hike upslope toward the prominent boulders for Stop 3.1. Stop 3.1. Ruby Dome Ranch Road, Lamoille Type Locality (40.69815°N, 115.48804°W) A broad, hummocky moraine at the mouth of Lamoille Canyon was designated the type locality for the Lamoille Glaciation by Sharp (1938), and is one of few locations in the Ruby and East Humboldt Ranges where moraines are preserved at the mountain front (Fig. 17). The lobate shape of the moraine suggests that it was formed by a piedmont glacier emanating from the mouth of Lamoille Canyon.

The Lamoille Glaciation has been correlated to the Illinoian Glaciation (MIS-6; 186–128 ka) based on weathering characteristics of moraines (Bevis, 1995; Wayne, 1984). However, cosmogenic 10Be surface-exposure dating of Lamoille-age moraines at Hennan Canyon, a few km south of Stop 3.1, yielded ages ranging from ca. 19 to 66 ka, suggesting a younger age for the Lamoille Glaciation (Briggs et al., 2004). We sampled tall, quartz-rich boulders atop the type Lamoille moraine at this site (Fig. 18) along with clasts in outwash gravel that grades from the moraine for cosmogenic 10Be surface-exposure dating. Moraine boulders display evidence of physical erosion at their surfaces; however, when combined with a depth-profile age of outwash gravel downvalley, cosmogenic exposure ages may provide more precise limits on the timing of the Lamoille Glaciation. Directions to Stop 3.2 Return to the vans and continue along the road into the Ruby Dome Ranch (Fig. 17). At the house, turn left and park at the gate. Leave the vehicles, pass through the gate, being sure to close it behind you, and hike south (right) along the two-track toward Seitz Canyon. At a stream gaging station near the canyon mouth (~0.9 mi) the two-track turns left and climbs up the canyon to the proximal slope of the prominent Lamoille-age right lateral moraine. Continue 0.75 mi to the Angel Lake moraine complex for Stop 3.2. Stop 3.2. Seitz Canyon (40.38166°N, 115.50751°W) Seitz Canyon was occupied by one of the largest valley glaciers in the Ruby Mountains during the Lamoille and Angel Lake Glaciations. The moraine sequence in this canyon is very well preserved, displaying prominent left- and right-lateral moraines mapped as Lamoille-equivalent that extend to the mouth of the

Figure 16. (A) The trench excavated by a backhoe in the eastern branch of the V-shaped cuspate spit at (~1715 m asl) the Ventosa site viewed at stop 2.7 (view in A is to southwest). (B) Sandy gravel and gravelly sands that comprise the spit and typify shoreline deposits of Lake Clover elsewhere in the basin.

Pleistocene glacial and pluvial records in northeastern Nevada canyon where they are offset by a normal fault (Fig. 17). In the canyon, Angel Lake–age moraines are inset to the prominent Lamoille-age lateral moraines. In both Seitz and Hennan Canyons, the Angel Lake terminal moraine displays three prominent ridge crests (Fig. 19), suggesting repeated occupation of the moraines during the last glaciation. Similar moraine morphology is observed elsewhere in the Great Basin; for example, in the Wallowa Mountains where glacier maxima were dated at ca. 21 ka and ca. 17 ka (Licciardi et al., 2004). We have sampled moraine boulders atop all three ridges of the Angel Lake terminal moraine complex and atop recessional moraines farther upvalley in Seitz Canyon for cosmogenic 10Be surface-exposure dating. Directions to Stop 3.3 Hike back to the vehicles and drive back out to the main road. Turn right to enter Lamoille Canyon (Fig. 17). Signs near the boundary of the Humboldt-Toiyabe National Forest announce the start of the Lamoille Canyon Scenic Byway, which extends

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for 12 mi up to Roads End at an elevation of 2675 m near the head of the canyon. Following along the north and east side of Lamoille Creek through a deep U-shaped valley, this road is one of the most scenic drives in the Great Basin. A picnic area near the canyon mouth (fee charged) provides a sheltered spot for lunch along the rushing waters of Lamoille Creek, while other pull-outs with interpretive signs along the way highlight important aspects of the canyon’s scenery and natural history. Stop 3.3 is located 3 mi from the canyon mouth on the right. Park on the right in the prominent pull-out overlooking the hanging valley of the Right Hand Fork. Stop 3.3. Confluence of Lamoille Canyon and the Right Hand Fork (40.66192°N, 115.43883°W) An interpretive sign at this pull-out asks the question “What is a Glacier?” in front of a spectacular example of glacial geomorphology. Straight ahead to the south is the glacial valley of

Figure 17. Topographic map showing the lower reaches of Lamoille and Seitz Canyons (portions of the U.S. Geological Survey 7.5′ Lamoille and Noon Rock quadrangles), including Stops 3.1 and 3.2 (squares). Dashed lines indicate mapped moraine crests, and circles indicate locations of boulders sampled for cosmogenic 10Be surface-exposure dating.

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the Right Hand Fork which enters Lamoille Canyon as a hanging valley. During the Lamoille Glaciation, ice from both valleys met at this point and continued westward to the moraine position at the canyon mouth visited in Stop 3.1. Depending on light conditions, parallel downvalley-sloping benches interpreted as lateral moraines of Lamoille age are visible on the northern valley wall (behind you as you face the Right Hand Fork). During the more recent Angel Lake Glaciation, the glacier in Lamoille Canyon failed to reach this far down valley, and the glacier in the Right Hand Fork terminated as a piedmont lobe sprawled across the valley floor below the interpretive sign. The right lateral moraine of this glacier is an obvious ridge to the left rising nearly 100 m above the valley floor (Fig. 20). We have sampled boulders along the crest of this feature for 10Be surface-exposure dating and processing is under way. This lateral moraine wraps around as a fragmented terminal moraine visible as a series of lower relief ridges below the road, and an extensive outwash valley train graded to this moraine extends downvalley to the right as a sagebrush-covered flat. The prominent Angel Lake–age right lateral moraine represents a major impediment to Lamoille Creek. Today the creek descends through a narrow gorge squeezed between the moraine and the road just upstream from the interpretive sign. Above the moraine, the stream meanders through a low-gradient reach where significant aggradation has occurred. Exposures below the road reveal several meters of fine silt unconformably overlying rounded cobble gravel, suggesting slackwater deposition in an impoundment. Discovery of this exposure raised the exciting possibility that the slackwater sediments represented a lake that formed during the LGM when the Right Hand Fork Glacier occupied the right lateral moraine position. However, OSL dating of the basal silt overlying the cobble gravel returned an age of 2.83 ± 0.25 ka B.P., indicating that the lacustrine fill was deposited in the late

Figure 18. Large erratic boulder atop the type Lamoille moraine viewed at Stop 3.1 (view to the west). This boulder was sampled for cosmogenic 10Be surface-exposure dating.

Holocene. Perhaps beaver activity or slope processes produced a dam that temporarily held back the waters of Lamoille Creek. The suitability of this location for damming the stream was not overlooked by early settlers in this area; remnants of an artificial dam are visible just at the top of the cataract. However, projection of this elevation upstream reveals that the historic dam would not have impounded water to the elevation of the sampled sediments. Directions to Stop 3.4 Return to vehicles and continue upvalley. Turn right at the entrance to Thomas Canyon Campground and park straight ahead on the left. Take care not to block access to the campground host’s trailer, the registration board, or any of the campsites. Leaving the vehicles, walk into the campground across the bridge spanning a now-dry channel of Lamoille Creek. Continue straight ahead, crossing a bridge over the active channel, and head into the brush behind the outhouse on the left. Follow obvious game trails ~50 m east up to the crest of a bouldery ridge for Stop 3.4. Stop 3.4. Confluence of Lamoille Canyon and Thomas Canyon (40.64921°N, 115.40557°W) Rising up out of the brush just upstream from the Thomas Canyon Campground is a prominent moraine ridge studded with boulders. The genesis of this ridge is somewhat equivocal: it could be a right lateral moraine from a glacier entering the main valley from Thomas Canyon, it could be a medial moraine from a combined Thomas-Lamoille Canyon glacier as mapped by Bevis (1995), or it could be a terminal moraine from the Lamoille

Figure 19. View southward across the lower part of Seitz Canyon moraines showing four different moraine ridges representing two different glaciations. The highest, longest ridge is a left-lateral moraine produced during the Lamoille Glaciation. The inner three moraines represent the Last Glacial Maximum Angel Lake Glaciation, which apparently involved multiple oscillations of the ice margin to produce these separate moraine crests. Samples from all of these moraines are being processed for cosmogenic 10Be surface-exposure dating.

Pleistocene glacial and pluvial records in northeastern Nevada Canyon Glacier. The latter is appealing because there is no other candidate elsewhere in Lamoille Canyon for the terminal position of ice during the LGM. However, the morphology of the ridge, including the curvature of its upper section, seems to indicate that it was formed by ice exiting Thomas Canyon. Unfortunately, because the heads of both canyons are quite close to one another to the south, both glaciers were eroding the same rock types and there is no diagnostic erratic lithology that could be used to distinguish the glacier source area. This leaves a strange dilemma. Lamoille Canyon is the largest and best developed U-shaped valley in the Ruby Mountains. It hosted the longest glacier in the Great Basin during the penultimate Lamoille Glaciation. Yet, the position of the glacial terminus in this valley during the LGM Angel Lake Glaciation cannot be recognized. One explanation is that the moraine was entirely eroded away by subsequent meltwater activity and slope processes after the Angel Lake deglaciation. This explanation is somewhat undermined, though, by the tremendous preservation of the moraine ridges at the mouth of the Right Hand Fork. Another more likely explanation is that the moraine has been buried and obscured by focused deposition of post-glacial mass-wasting deposits. Numerous examples of debris fans are visible along the steep sides of the entire canyon, and Wayne (1984) mapped over a dozen of them between the canyon mouth and head. Perhaps the Angel Lake–age moraine is unlocatable because it was entombed within one of these debris fans. Interestingly, numerical modeling simultaneously applied to glaciers in Lamoille, Thomas, and the Right Hand Fork Canyons indicates that combinations of temperature and precipitation necessary for growth of the latter two glaciers would have driven the Lamoille Glacier to a position just upstream from the

Figure 20. The ice-proximal slope of the Angel Lake–age right-lateral moraine deposited by the glacier entering Lamoille Canyon from the Right Hand Fork (view to the east). The moraine at this location is more than 85 m high and features a sharp crest studded with large boulders. Several of these have been sampled for cosmogenic 10Be surface-exposure dating.

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Thomas Canyon Campground (Wendler and Laabs, 2008), suggesting that the ridge above the campground may actually be the terminal moraine from the Lamoille Canyon Glacier. Directions to Stop 3.5 Return to the vans, drive back out to the Lamoille Canyon Road, and turn right to continue upvalley. The moraine ridge visited at Stop 3.4 is briefly visible through the trees, after which the road progressively curves around to head more due south. In the vicinity of the Terraces Picnic Area (fee charged) ledges of striated bedrock are visible alongside the road. Processing of samples collected from these ledges for cosmogenic surface-exposure dating is under way and will hopefully provide some constraint on the timing of deglaciation in the upper valley. Near-vertical exposures feature striations overlain on a wavy erosional pattern that may represent P-forms generated by subglacial meltwater erosion. A major debris flow chute near the Terraces Picnic Area reactivated in June of 2010, locally burying the road in more than 5 m of sediment, and causing evacuation of the Thomas Canyon Campground for several days. Mobilization of sediment in this event was driven by rapid melting of a deep snowpack. Above the Terraces Picnic Area, the valley widens and views become more extensive as trees diminish in height. The road passes far below the hanging valley occupied by Island Lake on the right and ends at the “Roads End” parking lot. Park the vehicles here for Stop 3.5. Stop 3.5. Roads End (40.60432°N, 115.37582°W) From Roads End, the spectacular extent of the head of Lamoille Canyon is revealed (Fig. 21). The U-shaped cross section of the valley is plainly visible, and the high hanging valley of Island Lake is obvious to the left when looking downstream (north). A well-graded trail to Island Lake provides an easy 2-mi

Figure 21. Northward view of the U-shaped valley of Lamoille Canyon from below Liberty Pass. The road leading to the Roads End turnaround (Stop 3.5) is visible on the valley floor.

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climb up into this cirque. About halfway up, the trail passes a bench formed by till that is likely a lateral moraine. This feature is located just below the bridge over the cataract draining from the lake at an elevation of 2824 m, indicating a former ice thickness of ~200 m in this part of the valley. Roads End is the northern terminus of the Ruby Crest trail, which extends over 40 mi southward along the crest of the range to terminate in Harrison Pass near the pull-out where we stopped on Day 2 to view the Franklin Valley. The trail climbs southward from Roads End and leaves the Lamoille watershed at Liberty Pass, a narrow notch visible high to the south from the parking lot. The timing of the ultimate deglaciation in the headwaters of Lamoille Creek is poorly constrained; however, a bog-bottom date of 13,000 ± 900 14C years was reported from the vicinity of Roads End (Wayne, 1984). This date was obtained on bulk sediment, which undoubtedly contributes to the wide error factor and makes it difficult to interpret. Nonetheless, taking the date at face value and calibrating it to calendar years indicates that the upper part of Lamoille Canyon was deglaciated by 13,300–17,900 yr B.P. Directions Back to Logan, Utah From Roads End at the head of Lamoille Canyon, retrace the route back down to the canyon mouth. Continue out through the Lamoille-type moraines, and turn right on NV-227. Pass through the small town of Lamoille, and turn left at the end of town on Crossroads Lane. Turn right at the next section boundary on Clubine Road, and then follow this main road (gravel) as it zigzags generally northward along the range front. After ~17 mi from Lamoille, turn left on NV-229 (paved), and after 2 mi, turn right following signs for Starr Valley. Go ~10 more miles and turn right at the intersection in front of Starr Valley Community Hall. Follow this paved road ~8 mi to the Welcome interchange at I-80. Enter I-80 eastbound and retrace the route from Day 1 back to Logan. (Travel 35 mi east to exit 378. Follow NV-233 to the Utah line, and continue on U.S.-30 to I-84. Head east to I-15, and then exit to follow U.S.-30 to Logan.) ACKNOWLEDGMENTS Our work in northeastern Nevada has been supported by National Science Foundation grants NSF-P2C2-0902586 to Munroe, NSF-P2C2-0902472 to Laabs, NSF-BCS-0808861 to Munroe, and a Geological Society of America Quaternary Geology & Geomorphology Division Gladys W. Cole Memorial Research award to Laabs. Thanks to the Humboldt-Toiyabe National Forest and the Bureau of Land Management for permission to conduct this research on federal land. M. Badding, L. Best, M. Bigl, J. Johnson, C. Kruger, L. Luna, N. Taylor, and L. Wendler assisted in the field. REFERENCES CITED Antevs, E.V., 1948, The Great Basin, with emphasis on glacial and postglacial times; 3, Climatic changes and pre-white man: Bulletin of the University of Utah, v. 38, no. 20, p. 168–191.

Bartlein, P.J., Anderson, K.H., Anderson, P.M., Edwards, M.E., Mock, C.J., Thompson, R.S, Webb, R.S., Webb, T., III, and Whitlock, C., 1998, Paleoclimate simulations for North America over the past 21,000 years; features of the simulated climate and comparisons with paleoenvironmental data; Late Quaternary climates; data synthesis and model experiments: Quaternary Science Reviews, v. 17, no. 6-7, p. 549–585, doi:10.1016/ S0277-3791(98)00012-2. Benson, L.V., May, H.M., Antweiler, R.C., Brinton, T.I., Kashgarian, M., Smoot, J.P., and Lund, S.P., 1998, Continuous lake-sediment records of glaciation in the Sierra Nevada between 52,600 and 12,500 14C yr B.P: Quaternary Research, v. 50, no. 2, p. 113–127, doi:10.1006/qres.1998.1993. Benson, L., Kashgarian, M., Rye, R., Lund, S., Paillet, F., Smoot, J., Kester, C., Mensing, S., Meko, D., and Lindström, S., 2002, Holocene multidecadal and multicentennial droughts affecting Northern California and Nevada: Quaternary Science Reviews, v. 21, no. 4–6, p. 659–682, doi:10.1016/ S0277-3791(01)00048-8. Bevis, K.A., 1995, Reconstruction of late Pleistocene paleoclimatic characteristics in the Great Basin and adjacent areas [Ph.D. dissertation]: Corvallis, Oregon, USA, Oregon State University. Birkeland, P.W., Machette, M.N., and Haller, K.M., 1991, Soils as a tool for applied Quaternary geology: Salt Lake City, Utah, USA, Utah Geological Survey Miscellaneous Publication no. 91-3. Bischoff, J.L., and Cummins, K., 2001, Wisconsin glaciation of the Sierra Nevada (79,000–15,000 yr B.P.) as recorded by rock flour in sediments of Owens Lake, California: Quaternary Research, v. 55, no. 1, p. 14–24, doi:10.1006/qres.2000.2183. Blackwelder, E., 1931, Pleistocene glaciation in the Sierra Nevada and Basin Ranges: Geological Society of America Bulletin, v. 42, no. 4, p. 865–922. Blackwelder, E., 1934, Supplementary notes on Pleistocene glaciation in the Great Basin: Journal of the Washington Academy of Sciences, v. 24, no. 5, p. 217–222. Briggs, R.W., et al., 2004, Cosmogenic Be-10 ages of Angel Lake and Lamoille Moraines and late Pleistocene slip rate of the rangefront normal fault, Ruby Mountains, Basin and Range, Nevada: Eos (Transactions, American Geophysical Union), v. 85, no. 47, 2004 Fall Meeting Suppl., abstract G11A-0773. Burr, T.N., and Currey, D.R., 1988, The Stockton Bar, in the footsteps of G.K. Gilbert, in Machette, M.N., ed., Lake Bonneville and neotectonics of the eastern Basin and Range Province; guidebook for field trip twelve: Salt Lake City, Utah, USA, Utah Geological Survey Miscellaneous Publication no. 88-1. Currey, D.R., and Oviatt, C.G., 1985, Durations, average rates, and probable causes of Lake Bonneville expansions, stillstands, and contractions during the last deep-lake cycle, 32,000 to 10,000 years ago, in Kay, P.A., and Diaz, H.F., Problems of and prospects for predicting Great Salt Lake levels—Proceedings of a National Oceanic and Atmospheric Administration Conference, 26–28 March 1985: Salt Lake City, Utah, USA, University of Utah Center for Public Affairs and Administration, p. 9–24. Enzel, Y., Wells, S.G., and Lancaster, N., 2003, Late Pleistocene lakes along the Mojave River, Southeast California, in Enzel, Y., Wells, S.G., and Lancaster, N., eds., Paleoenvironments and paleohydrology of the Mojave and southern Great Basin deserts: Geological Society of America Special Paper 368, p. 61–77. Garcia, A.F., and Stokes, M., 2006, Late Pleistocene highstand and recession of a small, high-altitude pluvial lake, Jakes Valley, central Great Basin, USA: Quaternary Research, v. 65, no. 1, p. 179–186, doi:10.1016/j .yqres.2005.08.025. Gilbert, G.K., 1890, Lake Bonneville: Reston, Virginia, USA, U.S. Geological Survey Monograph 1, 438 p. Howard, K.A., 2000, Geologic map of the Lamoille Quadrangle, Elko County, Nevada: Reno, Nevada, USA, Nevada Bureau of Mines and Geology Map 125, scale 1:24,000 (http://www.nbmg.unr.edu/dox/m125plate.pdf). Jewell, P.W., 2010, River incision, circulation, and wind regime of Pleistocene Lake Bonneville, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 293, no. 1-2, p. 41–50, doi:10.1016/j.palaeo.2010.04.028. Kutzbach, J.E., 1987, Model simulations of the climatic patterns during the deglaciation of North America, in Ruddiman, W.F., and Wright, H.E., Jr., eds., North America and adjacent oceans during the last deglaciation: Boulder, Colorado, USA, Geological Society of America, Geology of North America, v. K-3, p. 425–446. Laabs, B.J., and Munroe, J.S., 2008, Glacial and pluvial records of the Angel Lake glaciation in northeastern Nevada and inferences of latest

Pleistocene glacial and pluvial records in northeastern Nevada Pleistocene climate: Geological Society of America Abstracts with Programs, v. 40, no. 6, p. 147. Laabs, B.J.C., and Munroe, J.S., 2009, Developing the Late Quaternary record of pluvial Lake Clover, Northern Great Basin, USA (abst.): Transactions, American Geophysical Union, Fall Meeting 2009, abstract #PP11D-1351. Laabs, B.J.C., Plummer, M.A., and Mickelson, D.M., 2006, Climate during the last glacial maximum in the Wasatch and southern Uinta Mountains inferred from glacier modeling; Quaternary landscape change and modern process in western North America: Geomorphology, v. 75, no. 3-4, p. 300–317, doi:10.1016/j.geomorph.2005.07.026. Laabs, B.J., Bash, E.B., Refsnider, K.A., Becker, R.A., Munroe, J.M., Mickelson, D.M., and Singer, B.S., 2007, Cosmogenic surface-exposure age limits for latest-Pleistocene glaciation and paleoclimatic inferences in the American Fork Canyon, Wasatch Mountains, Utah, U.S.A: Eos (Transactions, American Geophysical Union) v. 88, no. 52, 2007 Fall Meeting Supplement, abstract PP33B-1274. Licciardi, J.M., Clark, P.U., Brook, E.J., Elmore, D., and Sharma, P., 2004, Variable responses of Western U.S. glaciers during the last deglaciation: Geology, v. 32, no. 1, p. 81–84, doi:10.1130/G19868.1. Lillquist, K.D., 1994, Late Quaternary Lake Franklin; lacustrine chronology, coastal geomorphology, and hydro-isostatic deflection in Ruby Valley and northern Butte Valley, Nevada [Ph.D. dissertation]: Salt Lake City, Utah, USA, University of Utah. Lindstrom, S., 1990, Submerged tree stumps as indicators of mid-Holocene aridity in the Lake Tahoe region: Journal of California and Great Basin Anthropology, v. 12, no. 2, p. 146–157. Lips, E.W., Marchetti, D.W., and Gosse, J.C., 2005, Revised chronology of late Pleistocene glaciers, Wasatch Mountains, Utah: Geological Society of America Abstracts with Programs, v. 37, no. 7, p. 41. Louderback, L.A., and Rhode, D.E., 2009, 15,000 Years of vegetation change in the Bonneville basin; the Blue Lake pollen record: Quaternary Science Reviews, v. 28, no. 3-4, p. 308–326, doi:10.1016/j.quascirev.2008.09.027 (special theme: modern analogues in Quaternary palaeoglaciological reconstruction, p. 181–260). Madsen, D.B., and Currey, D.R., 1979, Late Quaternary glacial and vegetation changes, Little Cottonwood Canyon area, Wasatch Mountains, Utah: Quaternary Research, v. 12, no. 2, p. 254–270, doi:10.1016/0033 -5894(79)90061-9. Mifflin, M.D., and Wheat, M.M., 1979, Pluvial lakes and estimated pluvial climates of Nevada, United States: Reno, Nevada, USA, Nevada Bureau of Mines and Geology, Bulletin 94, 57 p. Munroe, J.S., and Laabs, B.J.C., 2009, A 7000-year Lacustrine Record from Angel Lake, Nevada (abst.): American Geophysical Union, Fall Meeting 2009, abstract no. PP23C-1419. Osborn, G., and Bevis, K., 2001, Glaciation in the Great Basin of the Western United States: Quaternary Science Reviews, v. 20, no. 13, p. 1377–1410, doi:10.1016/S0277-3791(01)00002-6. Oviatt, C.G., 1997, Lake Bonneville fluctuations and global climate change: Geology, v. 25, no. 2, p. 155–158, doi:10.1130/0091-7613 (1997)025<0155:LBFAGC>2.3.CO;2. Oviatt, C.G., Currey, D.R., and Sack, D., 1992, Radiocarbon chronology of Lake Bonneville, eastern Great Basin, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 99, no. 3-4, p. 225–241, doi:10.1016/0031-0182(92)90017-Y.

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Oviatt, C.G., Thompson, R.S., Kaufman, D.S., Bright, J., and Forester, R.M., 1999, Reinterpretation of the Burmester core, Bonneville Basin, Utah: Quaternary Research, v. 52, no. 2, p. 180–184, doi:10.1006/ qres.1999.2058. Oviatt, C.G., Madsen, D.B., and Schmitt, D.N., 2003, Late Pleistocene and early Holocene rivers and wetlands in the Bonneville Basin of western North America: Quaternary Research, v. 60, no. 2, p. 200–210, doi:10.1016/ S0033-5894(03)00084-X. Phillips, F.M., Zreda, M.G., Benson, L.V., Plummer, M.A., Elmore, D., and Sharma, P., 1996, Chronology for fluctuations in late Pleistocene Sierra Nevada glaciers and lakes: Science, v. 274, no. 5288, p. 749–751, doi:10.1126/science.274.5288.749. Plummer, M.A., and Phillips, F.M., 2003, A 2-D numerical model of snow/ ice energy balance and ice flow for paleoclimatic interpretation of glacial geomorphic features: Quaternary Science Reviews, v. 22, no. 14, p. 1389– 1406, doi:10.1016/S0277-3791(03)00081-7. Reheis, M.C., 1999, Extent of Pleistocene lakes in the western Great Basin: Reston, Virginia, USA, U.S. Geological Survey Miscellaneous Field Studies MF-2323. Reinemann, S.A., Porinchu, D.F., Bloom, A.M., Mark, B.G., and Box, J.E., 2009, A multi-proxy paleolimnological reconstruction of Holocene climate conditions in the Great Basin, United States: Quaternary Research, v. 72, no. 3, p. 347–358, doi:10.1016/j.yqres.2009.06.003. Richmond, G.M., 1986, Stratigraphy and correlation of glacial deposits of the Rocky Mountains, the Colorado Plateau and the ranges of the Great Basin; Quaternary glaciations in the Northern Hemisphere: Quaternary Science Reviews, v. 5, p. 99–127. Rosenberg, B.D., Bigl, M.F., Munroe, J.S., and Ryan, P.C., 2011, X-ray diffraction analysis of weathering patterns in high-elevation glacial, periglacial, and eolian sediments in northern Nevada and Utah: Geological Society of America Abstracts with Programs, v. 43, no. 1, p. 114. Russell, I.C., 1885, Geological history of Lake Lahontan, a Quaternary lake of northwestern Nevada: Reston, Virginia, USA, U.S. Geological Survey Monograph 11, 287 p. Sharp, R.P., 1938, Pleistocene glaciation in the Ruby–East Humboldt Range, northeastern Nevada with abstract in German by Kurt E. Lowenstein: Journal of Geomorphology, v. 1, no. 4, p. 296–323. Simpson, J.H., 1876, Report of explorations across the Great Basin of the Territory of Utah for a direct wagon-route from Camp Floyd to Genoa, in Carson Valley, in 1859, by Captain J. H. Simpson (1876): U.S. Government Printing Office, Washington, D.C., USA, 518 p. Thompson, R.S., 1992, Late Quaternary environments in Ruby Valley, Nevada: Quaternary Research, v. 37, no. 1, p. 1–15, doi:10.1016/ 0033-5894(92)90002-Z. Wayne, W.J., 1984, Glacial chronology of the Ruby Mountains–East Humboldt Range, Nevada: Quaternary Research, v. 21, no. 3, p. 286–303, doi:10.1016/0033-5894(84)90069-3. Wendler, L.C., and Laabs, B.J.C., 2008, Reconstructions of latest Pleistocene glaciers and paleoclimate in the Ruby–East Humboldt Mountains, Nevada, USA: Geological Society of America Abstracts with Programs, v. 41, no. 3, p. 25. MANUSCRIPT ACCEPTED BY THE SOCIETY 23 FEBRUARY 2011

Printed in the USA

The Geological Society of America Field Guide 21 2011

Timing, distribution, amount, and style of Cenozoic extension in the northern Great Basin Christopher D. Henry* Nevada Bureau of Mines and Geology, University of Nevada, Reno, Nevada 89557-0178, USA Allen J. McGrew* Department of Geology, University of Dayton, Dayton, Ohio 45469-2364, USA Joseph P. Colgan* U.S. Geological Survey, Menlo Park, California 94025, USA Arthur W. Snoke* Department of Geology and Geophysics, University of Wyoming, Laramie, Wyoming 82071-2000, USA Matthew E. Brueseke* Department of Geology, Kansas State University, Manhattan, Kansas 66506, USA

ABSTRACT This field trip examines contrasting lines of evidence bearing on the timing and structural style of Cenozoic (and perhaps late Mesozoic) extensional deformation in northeastern Nevada. Studies of metamorphic core complexes in this region report extension beginning in the early Cenozoic or even Late Cretaceous, peaking in the Eocene and Oligocene, and being largely over before the onset of “modern” Basin and Range extension in the middle Miocene. In contrast, studies based on low-temperature thermochronology and geologic mapping of Eocene and Miocene volcanic and sedimentary deposits report only minor, localized extension in the Eocene, no extension at all in the Oligocene and early Miocene, and major, regional extension in the middle Miocene. A wealth of thermochronologic and thermobarometric data indicate that the Ruby Mountains–East Humboldt Range metamorphic core complex (RMEH) underwent ~170 °C of cooling and 4 kbar of decompression between ca. 85 and ca. 50 Ma, and another 450 °C cooling and 4–5 kbar decompression between ca. 50 and ca. 21 Ma. These data require ~30 km of exhumation in at least two episodes, accommodated at least in part by Eocene to early Miocene displacement on the major west-dipping mylonitic zone and detachment fault bounding the RMEH on the west (the mylonitic zone may also have been active during an earlier phase of crustal extension). Meanwhile, Eocene paleovalleys containing 45–40 Ma ash-flow tuffs drained eastward from

*[email protected]; [email protected]; [email protected]; [email protected]; [email protected]. Henry, C.D., McGrew, A.J., Colgan, J.P., Snoke, A.W., and Brueseke, M.E., 2011, Timing, distribution, amount, and style of Cenozoic extension in the northern Great Basin, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 27–66, doi: 10.1130/2011.0021(02). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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Henry et al. northern Nevada to the Uinta Basin in Utah, and continuity of these paleovalleys and infilling tuffs across the region indicate little, if any deformation by faults during their deposition. Pre–45 Ma deformation is less constrained, but the absence of Cenozoic sedimentary deposits and mappable normal faults older than 45 Ma is also consistent with only minor (if any) brittle deformation. The presence of ≤1 km of late Eocene sedimentary—especially lacustrine—deposits and a low-angle angular unconformity between ca. 40 and 38 Ma rocks attest to an episode of normal faulting at ca. 40 Ma. Arguably the greatest conundrum is how much extension occurred between ca. 35 and 17 Ma. Major exhumation of the RMEH is interpreted to have taken place in the late Oligocene and early Miocene, but rocks of any kind deposited during this interval are scarce in northeastern Nevada and absent in the vicinity of the RMEH itself. In most places, no angular unconformity is present between late Eocene and middle Miocene rocks, indicating little or no tilting between the late Eocene and middle Miocene. Opinions among authors of this report differ, however, as to whether this indicates no extension during the same time interval. The one locality where Oligocene deposits have been documented is Copper Basin, where Oligocene (32.5– 29.5 Ma) conglomerates are ~500 m thick. The contact between Oligocene and Eocene rocks in Copper Basin is conformable, and the rocks are uniformly tilted ~25° NW, opposite to a normal fault system dipping ~35° SE. Middle Miocene rhyolite (ca. 16 Ma) rests nonconformably on the metamorphosed lower plate of this fault system and appears to rest on the tilted upper-plate rocks with angular unconformity, but the contact is not physically exposed. Different authors of this report interpret geologic relations in Copper Basin to indicate either (1) significant episodes of extension in the Eocene, Oligocene, and middle Miocene or (2) minor extension in the Eocene, uncertainty about the Oligocene, and major extension in the middle Miocene. An episode of major middle Miocene extension beginning at ca. 16–17 Ma is indicated by thick (up to 5 km) accumulations of sedimentary deposits in half-graben basins over most of northern Nevada, tilting and fanning of dips in the synextensional sedimentary deposits, and apatite fission-track and (U-Th)/He data from the southern Ruby Mountains and other ranges that indicate rapid middle Miocene cooling through near-surface temperatures (~120–40 °C). Opinions among authors of this report differ as to whether this period of extension was merely the last step in a long history of extensional faulting dating back at least to the Eocene, or whether it accounts for most of the Cenozoic deformation in northeastern Nevada. Since 10– 12 Ma, extension appears to have slowed greatly and been accommodated by highangle, relatively wide-spaced normal faults that give topographic form to the modern ranges. Despite the low present-day rate of extension, normal faults are active and have generated damaging earthquakes as recently as 2008.

INTRODUCTION Although crustal extension in the North American Cordillera has been studied for nearly 100 years, many aspects of it—spatial distribution, timing, amount, structural style, and causes—remain controversial (Wernicke, 1992; Dickinson, 2002). These controversies are exemplified by the area of the northern Great Basin surrounding the Ruby Mountains–East Humboldt Range metamorphic core complex (RMEH) in northeastern Nevada (Fig. 1). This region underwent a series of Paleozoic to Mesozoic contractional episodes that significantly thickened the crust and probably raised its elevation, followed by moderate to large-magnitude

extension in the Cenozoic. Thermochronologic and thermobarometric studies of the metamorphic rocks in the RMEH and analysis of some Cretaceous and Paleogene strata are interpreted to indicate that major cooling, exhumation, and extension began in the Eocene or even Late Cretaceous, soon after—or even during—contraction, and continued through the middle Cenozoic (e.g., Hodges and Walker, 1992; McGrew and Snee, 1994; Camilleri and Chamberlain, 1997; Snoke et al., 1997; McGrew et al., 2000; Wells and Hoisch, 2008; Druschke et al., 2009a, 2009b). In contrast, geologic mapping and dating of Eocene to Miocene sedimentary and volcanic rocks and more recent lowtemperature thermochronology studies are interpreted to indicate

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that early to middle Cenozoic extension was relatively minor and that major extension took place in the middle Miocene (Henry, 2008; Wallace et al., 2008; Colgan and Henry, 2009; Colgan et al., 2010). The influence on extension of gravitational collapse of thickened crust, magmatic heating and weakening of the lithosphere, and changes in boundary conditions are similarly debated (Sonder and Jones, 1999; Rahl et al., 2002; Henry, 2008; Colgan and Henry, 2009). The characteristics of extension are significant not only for scientific reasons but are important to the formation and exploration for ore deposits and for hazards (Seedorff, 1991; Muntean et al., 2004; Cline et al., 2005; dePolo et al., 2011). This field trip examines the evidence for the timing of extension and uplift, emphasizing three different aspects of the regional geology: (1) the deeply exhumed metamorphic rocks and structures of the East Humboldt Range and evidence for their P-T-t paths; (2) Eocene and Miocene sedimentary deposits at Clover Hill (East Humboldt Range) that record the unroofing of the core complex; and (3) Eocene, rare Oligocene, and Miocene volcanic and sedimentary rocks that record the regional paleogeographic and tectonic evolution of northeastern Nevada at sites including the Copper Mountains area of northern Elko County and the northern Pequop Mountains east of the RMEH–Wood Hills metamorphic terrain (Figs. 2, 3, 4). Fault and tilt relationships between Cenozoic rocks are examined at several locations, especially in two ash-flow tuff- and sediment-filled paleovalleys that cross the extended region. The authors and field trip leaders represent the spectrum of interpretations about Cenozoic extension. STRATIGRAPHY AND ROCK UNITS Precambrian Though the presence of early Precambrian rocks in the deep subsurface of northeastern Nevada has long been inferred on the basis of the isotope geochemistry of younger plutonic rocks (Farmer and DePaolo, 1983; Wooden et al., 1997; Wright and Wooden, 1991), the gneiss complex of Angel Lake in the northern part of the East Humboldt Range provides the only reported exposures of Precambrian rocks predating the Neoproterozoic McCoy Creek Group (Lush et al., 1988). Although the age and petrogenesis of these polymetamorphic, migmatitic rocks has recently emerged as a matter of debate (Premo et al., 2008; McGrew and Snoke, 2010; Premo et al., 2010), new mapping by McGrew in the summer of 2009 coordinated with new U-Pb SHRIMP zircon geochronology by Premo sheds light on this controversy as briefly summarized below. The gneiss complex of Angel Lake consists of an orthogneiss unit and two paragneiss units occupying the core of the Winchell Lake fold-nappe, a map-scale, southward-closing recumbent fold with a WNW-trending hingeline (Figs. 5A, 5B). The core of the Winchell Lake fold is occupied by a newly mapped unit of strongly migmatitic biotite schist and impure metapsammittic rock yielding a suite of Archean detrital zircons suggesting a latest Archean or Paleoproterozoic protolith age overprinted by a

previously unrecognized metamorphic event at ca. 1.7 Ga (Fig. 5C) (W.R. Premo, 2010, personal commun.). Folded around this “core paragneiss” is the migmatitic orthogneiss of Angel Lake, a distinctively striped biotite monzogranitic orthogneiss with an Early Proterozoic age of 2450 ± 5 Ma based on a newly collected, minimally migmatized sample (Fig. 5B) (W.R. Premo, 2010, personal commun.). Folded around both the core paragneiss and the orthogneiss is an outer quartzite-pelitic schist unit that yields a detrital zircon suite with a strong Grenville-age spike, indicating a Neoproterozoic or younger depositional age and a probable correlation with the McCoy Creek Group of the northeastern Great Basin (W.R. Premo, 2010, personal commun.). All three units also host widespread amphibolite and garnet amphibolite sheets interpreted as metamorphosed mafic dikes and sills (Fig. 5B). Original contact relationships between these units are obscured by younger metamorphic and deformational overprints, and integrated geochemical, geochronologic and field-based studies are continuing in an effort to unravel the protracted history of these important exposures, which provide a unique opportunity to investigate and constrain the Precambrian history of continent formation as well as the subsequent polyphase tectonic history. These rocks are the focus of Stop 1-3. Paleozoic to Early Mesozoic Pre-Mississippian strata in northeastern Nevada are traditionally divided into a western facies assemblage consisting mostly of deepwater argillite and chert with associated dark quartzite, greenstone and limestone (Ordovician to Devonian Valmy, Vinini, and Woodruff Formations) and an eastern facies assemblage consisting of a thick sequence of Neoproterozoic to Cambrian clastic strata overlain by 5–7 km of lower Paleozoic shelf carbonates terminating in uppermost Devonian to Lower Mississippian shale (Figs. 3, 4). The eastern facies is inferred to have been deposited on cratonic North American basement, but this contact is not exposed in northeastern Nevada, with the possible exception of an outcrop of gneiss in the core of the Winchell Lake fold near Angel Lake in the East Humboldt Range (Stop 1-3). Both the eastern and western assemblages are overlapped by a Mississippian clastic sequence and then a return to shelfcarbonate deposition extending into the Lower Triassic. The western facies was thrust eastward over the eastern facies initially during the Late Devonian to Mississippian Antler orogeny along the regional Roberts Mountains thrust (Roberts, 1964; Smith and Ketner, 1977). Locally, however, the basal thrust cuts and therefore post-dates strata as young as Triassic (Coats and Riva, 1983). Jurassic and Cretaceous Plutons and Sedimentary Rocks Jurassic and Cretaceous plutons are scattered through northeastern Nevada (Coats, 1987; Wright and Snoke, 1993; Barton, 1996; Mortensen et al., 2000). They are relatively localized near the surface but presumably are more abundant at depth,

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? South

2-2 2-3

Fig. 14 Stop 2-1

Coal Mine Canyon

Double Mtn

Elko

746

2-4

Cornwall Mtn

Tv

Wood Hills

Tm

Pz

Pz

Q

Nanny Creek

Oasis

Fig. 16 Stop 3-1

Windermere Hills

Tm

Pz

?

233

Tjr

80

41°

Q

Tv

Pilot Rang

e

Figure 2. Geologic map of northeastern Nevada (simplified from Coats, 1987), showing locations of more detailed maps around major stops. Blue dot near Wells is epicenter of 21 February 2008 Mw 6.0 earthquake, which resulted from northwest-oriented extension on a 40°-striking, 55° southeast-dipping normal fault (dePolo et al., 2011).

0

0

Andesite and dacite lavas

Ta

N

Volcanic and sedimentary rocks

Tv

Eocene

Sedimentary rocks

Tm

Q

Bull Run Ts basin

Paleovalley with known or inferred flow direction, dashed where approximate

Q allluvium Miocene-Pliocene

Quaternary

41°

Central

Tv

Owyhee Plateau

North

Copper Tjr Basin

Pz

ts

Pz

Rub yM

Tv

In d e p end Mts ence

Jarbidge

EH u Ra mbold nge t

116°

Pequop M ts

42°

Cenozoic extension in the northern Great Basin 31

32

Henry et al.

Intrusive rocks

Volcanic rocks

Sedimentary rocks

PLIOCENE to HOLOCENE

Qal

Jarbidge Rhyolite

MIOCENE

Th

Indian Well Formation

Tjr

CRETACEOUS

Humboldt Formation (16–10 Ma) Thb (Breccia blocks)

Multiple Ash-flow tuffs

OLIGOCENE

EOCENE

Includes Elko Fm. shale, limestone conglomerate

No Paleocene rocks mapped in area of Figs. 2 or 4

JURASSIC TRIASSIC

Tgr

Upper Pz miogeocline

PERMIAN

High-angle fault Detachment fault

PENNSYLVANIAN

Mylonite zone Fold axis

MISSISSIPPIAN

9

Jurassic & Cretaceous intrusive rocks

Ppc

Park City Group (limestone & dolomite)

Pa

Arcturus Group (Pequop Fm) (limestone & sandstone)

e

8

CAMBRIAN

PROTEROZOIC

Approximate Thickness (km)

ORDOVICIAN

Neoproterozoic and Lower Paleozoic miogeocline

SILURIAN

7 Metamorphosed Proterozoic and Paleozoic sedimentary rocks

DEVONIAN

Roberts Mountains allochthon

K-Ar biotite age contour (Ma)

Mdc Mpj

Overturned fold

6

5

2

1

0

Ely Limestone Diamond Peak Formation & Chainman Shale

Ddg

Pilot Shale & Joana Limestone Guilmette Limestone

Dn

Nevada Formation (dolomite)

Slm

Lone Mountain Dolomite

Oe Op

Eureka Quartzite Pogonip Group (limestone)

u

Upper Cambrian shale & limestone

m

Middle Cambrian shale & limestone

pm

Prospect Mountain Quartzite (Lower Cambrian to latest Proterozoic)

4

3

Harrison Pass pluton

Mzgr

10

MAP SYMBOLS (Fig. 4)

25

Quaternary deposits

p s

McCoy Creek Group (Neoproterozoic)

?

Figure 3. Guide to stratigraphic units in the Ruby Mountains area of northeastern Nevada and explanation for Figure 4 (modified from Colgan et al., 2010). Right column is scaled to approximate thickness of Paleozoic to Triassic shelf sequence in the Ruby Mountains. Many Paleozoic units have different names in different mountain ranges.

DIAMOND

SULPHUR

Mtn.

RANGE

RIN GS

SP

VALLEY

GS

RANGE

RUBY

M

HUN TING TON

RU BY

MO

VAL LEY

e

r Rid g

Ceda

RA NG

30

UN 25 TA INS

PIÑON

RANGE

Carlin VALLE Y

20

LA MO I

Secret Pass

IC ED

CLOVER

LL

E

AD

RA NGE

VA LL EY

O BE

VALLEY

30

25

116º 115º

Fig. 10 Stops 1-1,Wells 1-2, 1-3

Clover Hill

E

Biotite K-Ar age contours

Delcer Buttes

INE

RA N

40º

115º

Figure 4. Geologic map of the Ruby Mountains area (from Crafford, 2007, and Colgan et al., 2010). Day 1 of field trip focuses on geology of the Ruby Mountains–East Humboldt Range metamorphic core complex around Clover Hill, the East Humboldt Range, and Secret Pass.

Mtns

MA VE RIC K

SPRIN

Bald

Butte

RANGE

10 km

DIAMOND

116º E G

40º VALLEY

CO RTE Z

Fig. 13 Stop 1-4 GE RAN

Elko

HUMBOLDT

E

41º EAST

PIN

Cenozoic extension in the northern Great Basin 33

C

A

0

F3 F2

F3

F1

0.5 km

D

B

34 Henry et al.

Cenozoic extension in the northern Great Basin as indicated by the abundance of granitic rocks in the RMEH (Snoke et al., 1997; McGrew et al., 2000; Howard, 2003; Lee et al., 2003). Late Cretaceous peraluminous granites formed by crustal anatexis in thickened crust are common (Miller et al., 1990; Lee et al., 2003), and exposures of Late Cretaceous migmatites in the deep-crustal rocks of the northern East Humboldt Range provide insight into how those granites may have formed (McGrew et al., 2000; Premo et al., 2008; McGrew and Snoke, 2010; Premo et al., 2010; Hallett and Spear, 2010). Cretaceous clastic rocks have been mapped in the eastern Cortez Range and at scattered localities throughout the Piñon Range (Fig. 3; Smith and Ketner, 1978). The only directly dated Cretaceous section, however, is the ~600-m-thick Newark Canyon Formation in the Cortez Range, known to be Early Cretaceous from fossils (Smith and Ketner, 1978) and a recent U-Pb zircon date of 116 ± 3 Ma on an interbedded tuff (Druschke et al., 2008). Eocene Sedimentary and Volcanic Rocks Two major sequences of volcanic and sedimentary rocks dominate the Cenozoic section of northeastern Nevada, Eocene and middle Miocene (Figs. 2, 3, 4). One area of Oligocene sedimentary rocks is known and may be particularly critical to understanding extension. The Eocene sequence includes conglomerate, sandstone, limestone, shale, andesite-dacite lava (and intrusive

Figure 5. Photographs of key rock types and field relationships in the northern East Humboldt Range. (A) Photograph, view to northwest, and sketch of the hinge zone of the Winchell Lake fold-nappe as exposed along the back wall of Winchell Lake cirque on the eastern face of the northern East Humboldt Range (Snoke et al., 1997). The core of the fold at this locality consists of dark outcrops of rusty-weathering graphitic paragneiss (rgs) surrounded by white-weathering Cambrian and Ordovician marble (DCmu) and Ordovician metaquartzite (DCmq). Enveloping the fold is a thick sequence of Neoproterozoic and Lower Cambrian flaggy quartzite and schist (CZqs). These metasedimentary rocks, particularly the metaclastic units, contain abundant sheet-like bodies of leucogranite orthogneiss. The hinge zone of this major structure trends west-northwest, approximately parallel to mineral elongation lineations in the deformed rocks. Closure is to the south. To the north, the fold is cored by near-Archean orthogneiss (Angel Lake orthogneiss) and associated paragneiss. A pre-folding, low-angle fault is inferred to separate the metasedimentary rocks in this photograph from the basement complex enclosed in the core of the fold. Scale is approximate. (B) Striped gray biotite monzogranitic orthogneiss of Angel Lake (Neoarchean(?) to earliest Paleoproterozoic) intruded by an orthoamphibolite sheet of probable Proterozoic age and by millimeter-scale to meter-scale seams, veins, and lenses of probable Late Cretaceous leucogranite as discussed in text. View to northwest. (C) Refolded fold developed in intricately interdigitated Neoarchean(?) orthogneiss and migmatitic Paleoproterozoic paragneiss. Note that three overprinted fold phases are indicated. View is ~2.5 m wide. (D) Photograph, view to west, and sketch of road-cut exposure of a low-angle normal fault exposed on the west flank of Clover Hill separating steeply dipping fanglomerate of the Miocene Humboldt Formation from a footwall of brecciated, strongly foliated calcite marble. A distinct red gouge lies along the fault (Mueller and Snoke, 1993a).

35

equivalents), and rhyolite to dacite ash-flow tuff (Ressel and Henry, 2006; Henry, 2008; Cook and Brueseke, 2010). The sedimentary rocks are exposed in small, widely scattered outcrops that are generally termed the Elko Formation because the largest and thickest exposures occur in the Elko Hills (Smith and Ketner, 1978). The Elko Formation in the Elko Hills consists of ~200 m of basal conglomerate and conglomeratic sandstone overlain by ~600 m of lacustrine shale, oil shale, claystone, siltstone, and minor water-laid tuff (Solomon et al., 1979; Ketner and Alpha, 1992; Haynes, 2003). Zircon U-Pb ages on water-laid tuffs in the basal conglomerate and near the top of the lacustrine sequence are 46.1 ± 0.2 Ma and 38.9 ± 0.3 Ma, respectively (Haynes et al., 2002; Haynes, 2003). Sequences in other areas are lithologically similar but thinner (Henry, 2008). The Elko Formation is generally interpreted to have accumulated in an extensional basin or basins formed during Eocene extension, including slip on the detachment fault bounding the west side of the Ruby Mountains (Solomon et al., 1979; Mueller and Snoke, 1993a; Satarugsa and Johnson, 2000; Haynes et al., 2002; Haynes, 2003). However, Henry (2008) interpreted that most Eocene sedimentary rocks accumulated in paleovalleys, probably where minor extension generated small half-graben basins. Cenozoic magmatism began at ca. 45 Ma in northeastern Nevada, generally post-dating the oldest sedimentary deposits (whose age is constrained to post-Paleozoic, and locally postTriassic), and was part of a southward-migrating belt that swept from Washington and Idaho, through northeastern Nevada and northwestern Utah, and into central Nevada in the Oligocene (Best and Christiansen, 1991; Christiansen and Yeats, 1992; Brooks et al., 1995a, 1995b; Humphreys, 1995; Henry and Ressel, 2000). Eocene magmatism in northeastern Nevada was dominated by andesitic to dacitic lavas and compositionally similar intrusions in numerous centers. The lavas are only locally interbedded with sedimentary deposits in paleovalleys. Major ash-flow tuffs to be examined on this field trip are the “45 Ma tuff,” one or more ca. 41–43 Ma plagioclase-biotite tuffs, the ca. 41.0 Ma tuff of Coal Mine Canyon, and the 40.0 Ma tuff of Big Cottonwood Canyon (Henry, 2008). The tuff of Big Cottonwood Canyon erupted from a caldera in the Tuscarora volcanic field (Fig. 2; Henry et al., 1999; Henry, 2008). Source calderas for the other tuffs were probably also in the Tuscarora and Bull Run basin areas, where the tuffs are thickest and most widely distributed, but have not been positively identified. Within ~20 km of known or probable sources, the tuffs apparently blanketed the area. At greater distances, the tuffs largely flowed through and were deposited in paleovalleys (Fig. 2). Volcanism largely ceased in Elko County between ca. 35 and 16 Ma. Abundant silicic intrusions in the Ruby Mountains and East Humboldt Range are as young as 29 Ma, however, indicating that magmatism continued in the subsurface (Wright and Snoke, 1993; MacCready et al., 1997; Howard, 2000; Miller and Snoke, 2009). The only known locally derived volcanic rock between those times is an apparently minor 31 Ma ash-flow tuff in the Piñon Range west of the Ruby Mountains. Oligocene, probably

36

Henry et al.

distal pyroclastic-fall tuffs are present in Copper Basin (Day 2; Table 1; McGrew and Foland, 2004). Miocene and Younger Sedimentary and Volcanic Rocks Miocene and younger clastic sedimentary rocks are abundant throughout northeastern Nevada. In the Elko area, the Miocene rocks are called the Humboldt Formation, which was originally defined by Sharp (1939) to include Eocene rocks, and later restricted to the Miocene deposits by Smith and Ketner (1976, 1978). The Humboldt Formation and age-equivalent units range from ca. 16.5 to 10 Ma everywhere they have been dated by 40 Ar/39Ar and tephrachronology (John et al., 2000; Wallace et al., 2008; Colgan et al., 2008; Colgan et al., 2010). These deposits include fine-grained fluvial and lacustrine deposits, primary and reworked tephra beds up to several meters thick, conglomerate whose clast composition depends on the local basement lithology, and very large (10s of meters) breccia blocks most often consisting of Paleozoic carbonate. Rhyolite lava flows are locally interbedded with the Humboldt Formation, although they are not considered part of the unit as typically defined. The Humboldt Formation at Stops 1-1 and 1-2 exhibits the types of deposits common to the unit. The Humboldt Formation was deposited in the hanging-wall basins of major Miocene normal faults during an episode of regional Basin and Range extension in the middle Miocene (Sharp, 1939; Smith and Ketner, 1978; John et al., 2000; Wallace et al., 2008; Colgan and Henry, 2009; Colgan et al., 2010). Late Miocene and younger (mostly Pliocene) deposits are rarely exposed but have been intercepted by drill holes in the deeper valleys and are lithologically similar to the Humboldt Formation. Volcanism resumed in northeastern Nevada ca. 16.5 Ma, coincident with major extension, and was primarily bimodal, basalt and rhyolite (Zoback et al., 1994; John et al., 2000; Colgan and Henry, 2009). The oldest rocks are minor basalt with Steens-type petrographic (abundant plagioclase phenocrysts to 5 cm long) and chemical characteristics near Copper Basin dated at 16.5 Ma (Rahl et al., 2002; Coats, 1964). Mafic rocks make up most of the Tv unit in the western and northern part of Figure 2 (Owyhee Plateau); published K-Ar ages range from ca. 16 to 8 Ma. Recent mapping, geochemistry, and 40Ar/39Ar geochronology indicates that most, if not all, of the basalt flows and shield volcanoes exposed on the Owyhee Plateau are younger than 11 Ma and are more primitive than the mid-Miocene flood-basalt related lava flows (e.g., they are high-alumina olivine tholeiite, Snake River olivine tholeiite, and transitional; Shoemaker, 2004). A voluminous group of distinctive, coarsely porphyritic rhyolite lavas, the Jarbidge Rhyolite, is widespread across northern Elko County (Fig. 2), but a few flows are also present northeast and southwest of Wells. The oldest dates from Jarbidge Rhyolite are 16.3 ± 0.03 Ma on a lava dome west of Copper Basin (M.E. Brueseke, unpublished 40Ar/39Ar data) and 16.15 ± 0.02 Ma on a flow south of Bull Run basin (C.D. Henry, unpublished 40Ar/39Ar data). Jarbidge Rhyolite lava flows just north of Copper Basin

are as young as 15.8 ± 0.06 Ma (M.E. Brueseke, unpublished Ar/39Ar data). Jarbidge-type rhyolites in the northeastern East Humboldt Range are referred to as the Willow Creek rhyolite complex, which yielded K-Ar (sanidine) ages in the interval 14.8–13.4 Ma (Mueller and Snoke, 1993b). The oldest Willow Creek rhyolite forms a sequence of flows. Rhyolite porphyry (13.8 ± 0.5 Ma) intruded the flows, and even younger “purple rhyolite” (13.4 ± 0.5 Ma) overlies volcaniclastic breccia in one fault-bounded block near the front of the range (Willow Creek rocks can be viewed on the way to Stop 1-1). 40

STRUCTURAL-TECTONIC-TOPOGRAPHIC EVOLUTION Paleozoic-Mesozoic Contraction and Crustal Thickening and Paleogene Erosion Northeastern Nevada underwent multiple episodes of contraction from the Late Devonian–Early Mississippian Antler orogeny through the Late Cretaceous Sevier orogeny (Roberts et al., 1958; Armstrong, 1968; Thorman, 1970; Miller et al., 1992; Poole et al., 1992; Camilleri et al., 1997; Taylor et al., 2000; DeCelles, 2004; Dickinson, 2006). Eastern Nevada lay in the hinterland of the Sevier orogenic belt (Armstrong, 1968; Miller and Gans, 1989; Camilleri et al., 1997; Vandervoort and Schmitt, 1990; Wright and Snoke, 1993; Howard, 2003; DeCelles, 2004), where the crust was inferred to have reached a thickness of 50– 60 km by end of the Cretaceous, based on restoration of Tertiary extension (Coney and Harms, 1984; Gans, 1987; DeCelles, 2004; DeCelles and Coogan, 2006), estimates of shortening in the overthrust belt (Thorman et al., 1991; Camilleri et al., 1997), and metamorphic mineral assemblages (Camilleri et al., 1997; McGrew et al., 2000; Lee et al., 2003). Northeastern Nevada was an eroding highland (Nevadaplano) at the beginning of the Cenozoic (Armstrong, 1968; Coney and Harms, 1984; Christiansen and Yeats, 1992; Dilek and Moores, 1999; DeCelles, 2004; Best et al., 2009), from which sediments were most likely carried eastward to the Uinta and Green River Basins (Baars et al., 1988; Hintze, 1993; Goldstrand, 1994; Davis et al., 2009), although there is little sedimentary record in Nevada until the Eocene (Fouch et al., 1979; Solomon et al., 1979; Vandervoort and Schmitt, 1990). Interpretations of the absolute surface elevation, timing of surface uplift, and paleotopography of northeastern Nevada during the Paleocene and Eocene vary widely, however. Coney and Harms (1984), Dilek and Moores (1999), DeCelles (2004), DeCelles and Coogan (2006), Best et al. (2009), and Cassel et al. (2010)—drawing analogies to present-day Tibet and the Andean Plateau—inferred surface elevations of >3 km, probably as a result of crustal thickening during the Cretaceous Sevier orogeny. Absolute Eocene elevations interpreted from fossil leaves in Copper Basin (Stop 2-1; present-day elevation 2.2 km) vary from 1.1 km (Axelrod, 1966a, 1966b), to 2.0 ± 0.2 km (Wolfe et al., 1998), to 1.6 ± 1.6 km or 2.8 ± 1.8 km (Chase et al., 1998). Mix et al. (2011)

Mineral

Biotite

Biotite

Biotite Biotite Biotite

" Biotite

Sanidine

Sanidine

050721-11

960702-1A 050719-4 050727-1

" 050721-13

H06-121

H06-123

Windermere Hills H06-108 Sanidine

Biotite

050720-3B

Dead Horse Formation 970728-3 Sanidine

050720-9C

Meadow Fork Formation 010801-5 Hornblende " Hornblende " Biotite " Biotite 050720-10 Biotite

Copper Basin 990704-1 Plagioclase

Sample

40

39

Reworked tuff (same as WC-15, Mueller et al., 1999)

Fine, white, pyroclasticfall tuff, uppermost White, pyroclastic-fall tuff, directly beneath flora locality White, biotite-rich, pumiceous pyroclasticfall tuff Ash-flow tuff Andesite, lower White pumiceous tuff, lower " White, biotite-rich, pumiceous pyroclasticfall tuff, lower Tuff of Big Cottonwood Canyon 45 Ma tuff

Pyroclastic-fall tuff " " " Fine pyroclastic-fall tuff, lower Very fine pyroclastic-fall tuff, basal

Plagioclase-phyric, 76 Creek Basalt (Steenstype)

–114.64339

–115.48560

–115.49297

626514 627724

626216 625235 626514

627385

626511

627290

625654

625425 625425 625425 625425 625629

628060

41.25447

41.74391

41.74840

4621344 4623975

4621502 4621654 4621344

4624325

4623195

4624760

4623690

4624225 4624225 4624225 4624225 4624013

4623835

34.75 ± 0.12

44.90 ± 0.07

40.02 ± 0.10

43.4 42.6

40.0 42.7 42.1

38.2

39.6

37.9

31.8

28.3 28.2 27.0 28.6 31.0

16.3

47.3 ± 0.2

41.3 ± 0.1 42.7 ± 0.1

41.5 ± 0.2

39.8 ± 0.2

37.4 ± 0.2

29.2 ± 0.2 29.7 ± 0.3 29.3 ± 0.4 31.4 ± 0.3 32.5 ± 0.2

16.5 ± 0.2

43.8 ± 0.9 46.3 ± 1.0

40.9 ± 0.2 42.5 ± 0.2 42.0 ± 1.7

40.0 ± 1.3

39.6 ± 0.2

37.8 ± 0.2

29.4 ± 0.4 29.5 ± 0.4 27.8 ± 1.0 29.3 ± 1.3 33.7 ± 0.3

16.3 ± 0.1

TABLE 1. Ar/ Ar AGES, COPPER BASIN AND WINDERMERE HILLS, NEVADA Rock type and part of unit Easting or Northing or Ages longitude latitude (Ma) Weighted Integrated Plateau Isochron mean

255 ± 3 302 ± 6

302 ± 3 309 ± 9 258 ± 5

303 ± 5

305 ± 6

284 ± 10

293 ± 2 294 ± 2 302 ± 5 304 ± 5 271 ± 4

298 ± 2

36

Ar/ Ari

40

C.D. Henry (unpublished)

Henry (2008)

Henry (2008)

McGrew and Foland (unpub.) McGrew and Foland (unpub.)

Rahl et al. (2002) McGrew and Foland (unpub.) McGrew and Foland (unpub.)

McGrew and Foland (unpub.)

McGrew and Foland (unpub.)

Rahl et al. (2002)

McGrew and Foland (unpub.)

McGrew and Foland (unpub.) McGrew and Foland (unpub.) McGrew and Foland (unpub.) McGrew and Foland (unpub.) McGrew and Foland (unpub.)

Rahl et al. (2002)

Reference

Cenozoic extension in the northern Great Basin 37

38

Henry et al.

interpreted stable-isotope data to indicate uplift migrated southward through western North America during the early Cenozoic and Eocene elevations of ~3.4 km in the Elko and Copper basins. Major Eocene and older(?) paleovalley systems that drained eastward across northeastern Nevada record erosion of the Nevadaplano (Figs. 1, 2; Henry, 2008). Paleovalleys are discontinuously exposed across a series of ranges that formed predominantly during middle Miocene or later Basin and Range extension. Recognition and correlation of paleovalleys is based on (1) paleovalley geometry where ash-flow tuffs and sedimentary rocks crop out within valleys incised into Paleozoic rocks, (2) recognition of depositional contacts of paleovalley fill against walls, (3) wedging out of paleovalley fill against paleovalley walls, (4) continuity of individual tuffs and distinctive groups of tuffs along the paleovalleys, and (5) the considerable thicknesses and dense welding of tuffs far from their source. The north paleovalley can be traced from the Bull Run basin through Cornwall Mountain to Copper Basin, where it is joined by a probable tributary from the southwest and becomes buried eastward beneath Jarbidge Rhyolite (Fig. 2; Henry, 2008; McGrew and Vance, 2008; Cook and Brueseke, 2010). Insufficient data are available to follow the paleovalley farther east. The “45 Ma tuff” crops out all along this paleovalley and may have erupted from a caldera at the western end north of Bull Run basin. A plagioclase-biotite tuff, tuff of Coal Mine Canyon, tuff of Big Cottonwood Canyon, and a moderately thick section of conglomerate, tuffaceous sandstone, and shale overlie the 45 Ma tuff in Copper Basin (Stop 2-1). The central paleovalley can be traced ~150 km from near the Tuscarora volcanic field eastward over the Independence Mountains as far as Nanny Creek (Stop 3-1; Pequop Mountains; Fig. 2). Although gaps exist across several modern basins, continuity is inferred based on proximity and trend of individual segments and similar stratigraphy within them, especially the presence of the tuff of Big Cottonwood Canyon. The tuff of Big Cottonwood Canyon was able to flow 150 km (present day) from Tuscarora to Nanny Creek. The amount of post–40 Ma extension across that distance is unknown, but, even if it was 100%, the tuff must have flowed 75 km. What may be a branch of the central paleovalley passes through the Windermere Hills, where Mueller et al. (1999) interpreted Eocene sedimentary rocks to fill a half-graben above an east-rooted, Eocene low-angle detachment fault. Henry (2008) interpreted the same deposits to fill an east-draining paleovalley, consistent with cobble imbrications found by Mueller (1992), but probably dammed during small-magnitude Eocene extension. Eocene volcanic rocks, the sedimentary section, and overlying middle Miocene deposits are all tilted the same amount, which suggests major extension was middle Miocene or younger. Eocene (and Oligocene?) deposits on both west and east sides of the Pilot Range form a thick section of tuffaceous sedimentary rocks, possibly some primary ash-flow tuff, and coarse conglomerate (Miller, 1985; Miller et al., 1993; this study) that may be in the eastern continuation of this paleovalley.

A south paleovalley extends from the southern Independence Mountains through Coal Mine Canyon and may connect eastward across the northern East Humboldt Range to Clover Hill, where a thick section of Eocene and Miocene sedimentary rocks is exposed. The relative proportions of each are uncertain, but a basal conglomerate and ash-flow tuff are almost certainly Eocene. A cobble in the basal conglomerate was dated at ca. 38 Ma (Brooks et al., 1995a, 1995b). The upper part of the section contains a ca. 15.5 Ma tephra that is overlain by Willow Creek rhyolite (Jarbidge Rhyolite; Snoke et al., 1997; J.P. Colgan, K.A. Howard, and C.D. Henry, unpublished data). Coarsely clastic and breccia-bearing sedimentary rocks make up most of the section. Stops 1-1 and 1-2 show the sedimentary section, structure of the Clover Hill area, and highly metamorphosed rocks of the adjacent RMEH. Figure 2 shows a possible paleovalley crossing the Ruby Mountains, based on the thick section of Elko Formation in the Elko Hills, a thick section of Miocene or Eocene deposits identified by seismic-reflection methods in the valley west of the Ruby Mountains (Satarugsa and Johnson, 2000), and a thick section of dacitic and andesitic lavas with a basal coarse conglomerate in the southern East Humboldt Range (Brooks et al., 1995a, 1995b; this study). Satarugsa and Johnson (2000) interpreted the subsurface rocks to be middle Miocene, but distinguishing Eocene and Miocene rocks without drill hole data is uncertain. Cenozoic Extension Interpreted episodes of extension began with initial exhumation of the Ruby Mountains between the Late Cretaceous and Eocene (McGrew and Snee, 1994; Camilleri and Chamberlain, 1997; McGrew et al., 2000) and development of the Elko basin of the Elko Hills at ca. 46 Ma (Haynes et al., 2002; Cline et al., 2005; Hickey et al., 2005). A major episode of extension began in the middle Miocene throughout northern Nevada (Zoback et al., 1994; Miller et al., 1999; John et al., 2000; Wallace et al., 2008; Colgan and Henry, 2009; Colgan et al., 2010), and extension continues to today (e.g., 2008 Mw 6.0 Wells earthquake; Fig. 2). How much extension occurred at different times in the Cenozoic is the focus of this trip. The following four subsections discuss different approaches and perspectives. Evidence from Thermochronology and Thermobarometry of the RMEH (A.J. McGrew) Integrated thermobarometric and thermochronologic results indicate ~170 °C of cooling and 4 kbar of decompression of the RMEH between ca. 85 Ma and ca. 50 Ma (e.g., Dallmeyer et al., 1986; Dokka et al., 1986; Hurlow et al., 1991, Hodges et al., 1992; McGrew and Snee, 1994; McGrew et al., 2000; Hallett and Spear, 2010) (Fig. 6). However, maximum tectonic burial of the RMEH occurred before 85 Ma and is recorded by peak pressure conditions represented by a relict kyanite + staurolite assemblage that is preserved locally at high structural levels in the northern East Humboldt Range and better preserved at the still-higher

Cenozoic extension in the northern Great Basin structural levels of nearby Clover Hill (Snoke, 1992). Based on Gibbs Method modeling of the Clover Hill assemblage, Hodges et al. (1992) interpreted isothermal decompression from 9– 10 kbar to 5.0–6.4 kbar at temperatures of 550–630 °C before partial degassing inferred from an 40Ar/39Ar hornblende age spectrum in the mid-Cretaceous (Dallmeyer et al., 1986). At the deeper structural levels of the northern East Humboldt Range, peak metamorphism recorded by growth of sillimanite occurred contemporaneously with in situ partial melting in the Late Cretaceous, resulting in re-equilibration of most mineral

39

assemblages, although growth zoning has been preserved in some garnets at higher structural levels (Hallett and Spear, 2010; McGrew et al., 2000; McGrew and Snoke, 2010). McGrew et al. (2000) documented a major phase of leucogranite generation at 84.8 ± 2.8 Ma synkinematic with fold-nappe emplacement and metamorphism above the second sillimanite isograd. More recently, U-Pb SHRIMP analyses of zircon rims from the orthogneiss of Angel Lake confirm the presence of abundant Late Cretaceous melt (yielding a spectrum of essentially concordant ages from 72 to 91 Ma) (Premo, et al., 2008; McGrew et al., 2009).

12.0

Metapelites (McGrew et al., 2000) 15°C/km

11.0

Amphibolite (McGrew et al., 2000) Pre-Late K

10.0

25°C/km 85 Ma – peak metamorphism – migmatization

9.0

Metapelitic mylonites, SW East Humboldt Range (Hurlow et al., 1991) Clover Hill Gibbs Results (Hodges et al., 1992) N. Ruby Mts Gibbs Results (Hodges et al., 1992)

Pressure (kbar)

8.0 Al2SiO5 Phase Relations (Bohlen et al., 1991) 7.0

Linear (Metapelites (McGrew et al, 2000)

6.0

40°C/km

5.0

40-50 Ma – 40Ar/39Ar hbl closure at shallow structural levels

25°C/km Reference geotherms

60°C/km

4.0

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Temperature (°C) Figure 6. Interpretative diagram of P-T results from the Ruby Mountains–East Humboldt Range metamorphic core complex (RMEH) from McGrew et al. (2000) combined with previously published results from elsewhere in the RMEH (Hurlow et al., 1991; Hodges et al., 1992). Error ellipses have been left off to simplify viewing. The thick, colored arrows show the general P-T-t path inferred for the East Humboldt Range from late Mesozoic time to ca. 15 Ma based on the integration of P-T data with thermochronometric and other constraints. The earliest phase in the P-T-t path, illustrated by the upper arrow, records initial tectonic burial to kyanite-grade conditions before overprinting in the Late Cretaceous by peak metamorphism above the second sillimanite isograd and widespread migmatization. The second arrow represents a linear fit to the P-T results of McGrew et al. (2000) with a slope of 0.025 kbar/°C interpreted as a decompressional cooling path based on a variety of decompressional reaction textures. This segment is bracketed between 85 Ma and 40Ar/39Ar hornblende cooling ages of 40–50 Ma (at shallow structural levels) or 30–40 Ma (at deep structural levels). Thinner arrows represent net Gibbs model paths for samples from Clover Hill and the northern Ruby Mountains (Hodges et al., 1992). The lowermost heavy arrow illustrates inferred cooling based on late Oligocene to early Miocene 40Ar/39Ar muscovite and biotite and fission-track apatite cooling ages (Dallmeyer et al., 1986; Dokka et al., 1986; McGrew and Snee, 1994) combined with the constraint that metamorphosed footwall clasts do not appear until deposition of the middle Miocene Humboldt Formation. Additionally, ages of 11.6–13.8 Ma on authigenic illite from fault gouge in Secret Pass are interpreted to record last major activity of the RMEH detachment and associated faults (Haines and van der Pluijm, 2010). The modern-day geothermal gradient in the Basin and Range province is ~25 °C/km, whereas the Battle Mountain heat-flow high is characterized by geothermal gradients as high as 40–75 °C/km.

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Collectively, P-T estimates based on mineral rim thermobarometry from the northern East Humboldt Range define a data array extending from 798 °C, 9.3 kbar to 630 °C, 5.2 kbar, inferred to represent equilibration at different points along a decompressional P-T-t path (Fig. 6) (McGrew et al., 2000). Including data interpreted to constrain the P-T conditions during mylonitization at higher structural levels in the southwestern part of the East Humboldt Range extends this trend down to ~550 °C, 3–5 kbar (Hurlow et al., 1991). 40 Ar/39Ar hornblende isochron ages ranging from 50 to 63 Ma at higher structural levels in the East Humboldt Range imply that much if not all of this decompressional P-T-t path occurred between the Late Cretaceous and early Eocene, although the lowest P-T results, inferred to represent conditions during extensional mylonitic deformation, may have equilibrated in the late Eocene (Fig. 6; Hurlow et al., 1991; McGrew et al., 2000). Aluminum-inhornblende barometry on late Eocene (40–36 Ma) quartz diorites exposed in the RMEH further supports this interpretation, indi-

cating that these granitic rocks intruded at mid-crustal levels and pressures of ~5.5 kbar, near the base of the P-T trend described above and illustrated in Figures 6 and 7 (Snoke et al., 2004). Evidence of extension at upper crustal levels during the Late Cretaceous to the Eocene (ca. 46 Ma) in the RMEH environs is scant, although Camilleri and Chamberlain (1997) interpret a major normal fault, the Pequop Fault, that they bracket within this time frame. However, deep-crustal spreading or diapiric upwelling triggered by partial melting of a buoyant, thermally weakened lower-crustal root could have transferred this terrain from deep-crustal to mid-crustal levels without greatly extending the overlying rigid upper crustal layer (e.g., Wernicke and Getty, 1997; McGrew, 2002; Hodges, 2006). MacCready et al. (1997) provided evidence of such channel flow from the northern Ruby Mountains. However, to maintain crustal-scale strain compatibility, any such model would require a decoupling zone at mid-crustal levels separating the ductilely flowing deeper crustal layer from the overlying, more rigid upper crust. In this context,

900

800 Late Eocene mylonitic recrystallization at 3 - 5 kbar (Fig. 6)

700 40Ar/39Ar Hornblende

Late K to Eocene decompression to < 6 kbar (Fig. 6)

cooling ages, 30–38 Ma, at deep structural levels

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Age (Ma) Figure 7. Generalized time-temperature curve for the northern East Humboldt Range synthesizing results modified from McGrew and Snee (1994), and summarizing results from other studies (Dallmeyer et al., 1986; Dokka et al., 1986; Haines and van der Pluijm, 2010). Vertical black bars represent the inferred closure temperature ranges for hornblende (480–580 °C), muscovite (350–400 °C), and biotite (250–300 °C), respectively. Pressure estimates for metamorphism above hornblende closure temperatures derive from thermobarometric studies summarized in Figure 6. Short-dash line is a lower bound for relatively shallow structural levels, and long-dash line is an upper bound for deeper structural levels. Important magmatic events are the intrusion of a thick sheet of hornblende biotite quartz diorite at 40 ± 3 Ma and injection of widespread biotite monzogranitic sheets at 29 ± 0.5 Ma (Wright and Snoke, 1993).

Cenozoic extension in the northern Great Basin large-scale fold-nappes such as the Winchell Lake fold-nappe in the East Humboldt Range or the Soldier Peak or Lamoille Canyon fold-nappes in the Ruby Mountains could be manifestations characteristic of the “roof zone” of a laterally spreading and upwelling deep-crustal migmatite complex as presaged by Howard (1980). In this light, it is interesting that the Winchell Lake fold-nappe is both temporally and spatially associated with the “roof” of the migmatite complex in the very zone where the rheologically critical melt percentage is likely to have been achieved (Vanderhaeghe et al., 2001). Regardless of the pre-Eocene history of the RMEH, a variety of lines of evidence indicate that it remained at mid-crustal depths and elevated temperatures at least into the Oligocene: (1) Aluminum-in-hornblende barometry on 40–36 Ma quartz diorites indicates that they were intruded at mid-crustal depths, ~5.5 kbar (Snoke et al., 2004). (2) Both the 36–40 Ma intrusive suite and a younger, 29 Ma biotite monzogranitic intrusive suite were foliated, mylonitized, and locally complexly folded after emplacement. Moreover, petrographic relationships such as dynamic recrystallization of biotite along shear bands and crystal-plastic deformation of feldspars in these rocks indicates that the plastic deformation occurred at elevated temperatures (probably at least 450 °C). (3) Hurlow et al. (1991) and Hodges et al. (1992) document P-T conditions of 550–650 °C, 3.0–4.3 kbar based on thermobarometry of syn-mylonitic mineral assemblages in pelitic schist from the RMEH mylonitic shear zone in the southwestern East Humboldt Range (Fig. 6). (4) Quartz crystallographic preferred orientations (CPOs) are interpreted to record WNW- or ESE-directed shearing within and beneath the mylonitic zone, and show the hallmarks of recrystallization at temperatures as high as 600–700 °C (e.g., strong c-axis maxima parallel to the Y-strain axis and/or crossed girdle fabrics with a 90° opening angle at deeper structural levels, resembling relationships in the Saxony granulite terrain) (McGrew and Casey, 1993; MacCready, 1996; Law et al., 2004). (5) Deep structural levels of the RMEH last cooled through 40 Ar/39Ar hornblende closure temperatures during the late Eocene to the Oligocene (25–36 Ma) (Figs. 6, 7) (Dallmeyer et al., 1986; McGrew and Snee, 1994). Cooling through 40Ar/39Ar biotite closure temperatures (250–300 °C) occurred diachronously across the RMEH from ESE to WNW over the interval ca. 36–21 Ma, giving rise to an asymmetrical, WNW-younging cooling age “chrontour” pattern across the core complex (Kistler et al., 1981; Dallmeyer et al., 1986; McGrew and Snee, 1994) (Figs. 4, 7). The cooling age chrontours are perpendicular to mylonitic stretching lineation associated with numerous WNW-directed normal-sense kinematic indicators, leading to the common interpretation that the chrontour pattern tracks the unroofing history of the lower plate as the mylonitic shear zone progressively translated the upper plate to the WNW. However, the pattern could alternatively be interpreted as a transect through an exhumed, ESE-rotated partial retention zone (e.g., Colgan et al., 2010).

41

Finally, based on apatite, zircon, and sphene fission-track results from the RMEH, Dokka et al. (1986) interpreted rapid cooling between ~285 °C and ~70 °C between ca. 25 and 23 Ma, which suggests that exhumation was largely complete by the early Miocene. Early Miocene cooling of the RMEH is also constrained by the fact that metamorphic foliations in both the mylonitic zone and the underlying infrastructure are cut by unmetamorphosed and undeformed amygdaloidal basaltic dikes with chilled margins dated ca. 17–15 Ma (Snoke, 1980; Hudec, 1992). Consequently, it would appear that mylonitization had ceased and the RMEH had cooled to temperatures below ~120 °C, corresponding to paleodepths less than 4–6 km depending on the geothermal gradient assumed, by the early Miocene. While internally consistent, this cooling history contrasts with a recently documented cooling history for the Secret Pass area and continued slip on the RMEH detachment fault to ca. 12 Ma (Figs. 6, 7) (Haines and van der Pluijm, 2010). In addition, the accumulation of up to 5 km of middle Miocene fill in the adjoining basins to the west and evidence for rapid middle Miocene extension and exhumation in the southern Ruby Mountains pose a further challenge to the idea that most extension was completed by ca. 20 Ma (see “Miocene Basins and Thermochronology” section) (Satarugsa and Johnson, 2000; Colgan et al., 2010). Evidence from Eocene Paleovalleys and Ash-Flow Tuffs (C.D. Henry) The (1) continuity of paleovalleys across northeastern Nevada in the Eocene (Figs. 1, 2), (2) ability of Eocene ash-flow tuffs to flow long distances in the paleovalleys, (3) presence of relatively thin (<1 km) Eocene (≥41 Ma and ca. 38–35 Ma) sedimentary/lacustrine deposits, and (4) mostly concordant contacts between Eocene and middle Miocene deposits suggest one or two episodes of minor Eocene extension and little if any extension between the Eocene and middle Miocene (Henry, 2008; Colgan and Henry, 2009). A fundamental assumption is that tilting is a proxy for extension. The first two observations preclude northeastern Nevada being either a series of north-striking basins and ranges similar to the present topography or a single large basin having overall low relief. The tuff of Big Cottonwood Canyon flowed at least 75 km from source to its most distal known outcrop (Nanny Creek, Stop 3-1; original, pre-extension distance), and the tuff’s thickness and dense welding there suggest that it traveled significantly farther. Ash-flow tuffs erupted onto any low-relief surface would have spread radially and dispersed, and therefore would not show the linear distribution that they do or traveled very far. If erupted into a lake, these tuffs would have formed water-laid tuffs instead of thick, densely welded deposits. Tuffs erupted into basin-and-range topography like today also would have been channelized, but into wide north-trending paleovalleys, and would not have been able to reach distant locations to the east. Although low-density pyroclastic flows can surmount major ridges far from source, such flows are so dilute that they do not generate thick, densely welded deposits (Woods et al., 1998).

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Continuous paleovalleys had been eroded to their full depth by 45–40 Ma—the age range of ash-flow tuffs found in them— but may have formed much earlier. The paleovalleys are deep (possibly up to 1.6 km) but wide (6–8 km) and had low-relief interfluves, which seem characteristic of relatively deeply eroded topography. Formation of the paleovalleys during regional surface uplift related to the Late Cretaceous Sevier orogeny is therefore plausible. The third and fourth observations support possibly two episodes of low-magnitude extension, one before ca. 41 Ma and possibly as old as 46 Ma and a second between ca. 40 and 38 Ma. The Eocene sedimentary—especially lacustrine—deposits seem to require deposition in extensional basins, not only along paleovalleys. The deposits are widespread over northeastern Nevada but as small, discontinuous packages. Lacustrine deposits fall into two age groups: older than ca. 40 Ma in the Elko Hills and at Double Mountain, Coal Mine Canyon, and Copper Basin (Haynes et al., 2002; Rahl et al., 2002; Haynes, 2003; Henry, 2008; this report) and ≤38–35 Ma in the Windermere Hills (Mueller et al., 1999). These deposits are ~600 m thick in the Elko Hills and a few hundreds of meters thick elsewhere (Solomon et al., 1979; Mueller et al., 1999; Haynes, 2003; Henry, 2008). The only definite angular unconformity between rocks in the age range 45–16 Ma in most of northeastern Nevada is a ≤15° unconformity between the Elko Formation and 38 Ma volcanic rocks (Smith and Ketner, 1976; Solomon et al., 1979; Brooks et al., 1995a; Henry and Faulds, 1999; Haynes, 2003; Cline et al., 2005), immediately preceding the second episode of Eocene deposition. Where adequately studied, sub-unconformity rocks were tilted to the southeast. Southeast-tilted, pre–38 Ma rocks are present elsewhere but not overlain by younger rocks. For example, 41–40 Ma volcanic rocks in the Independence Mountains were also tilted to the southeast, but younger Eocene rocks are absent there (Muntean and Henry, 2007; Henry, 2008). The 46–39 Ma Elko Formation in the Elko Hills shows no fanning of dips (Haynes, 2003), which indicates no tilting and probably relatively little extension during its deposition. Similarly, Eocene and Miocene deposits are concordant in most areas of northeastern Nevada (Best and Christiansen, 1991; Thorman et al., 1991; Mueller, 1993; Brooks et al., 1995a; Mueller et al., 1999; Henry, 2008), which precludes significant tilting and extension. The distribution, characteristics, and timing of Eocene sedimentary and lacustrine deposits are consistent with their deposition in small depocenters along paleovalleys where they were intersected by Eocene high-angle normal faults. An ~65°-dipping normal fault with ~1 km of slip could form a half-graben basin along a paleovalley that would accumulate ~1 km of sediment. However, the thinness of Eocene deposits (≤1 km compared to as much as 5 km of middle Miocene and younger basin fill; Satarugsa and Johnson, 2000) and the small degree of angular unconformity imply small-magnitude extension in the Eocene. The absence of sedimentary deposits between ca. 38 and 16 Ma over most of northeastern Nevada, with the significant exception of Copper Basin, indicates no depocenters existed to

accumulate sediments. This observation implies little or no tilting and formation of half-graben basins, which suggests little if any extension in the upper crust. Whether or not the paleovalleys continued to be active drainages is uncertain, because no deposits of this age range are found in them. However, deposition continued throughout the Oligocene in the Uinta, Flagstaff, and Claron Basins in Utah (Baars et al., 1988; Hintze, 1993; Bryant et al., 1989; Davis et al., 2009), although whether northeastern Nevada was a source area is unknown. Eocene and Oligocene Extension: Copper Basin (A.J. McGrew) The Copper Mountains and adjacent Copper Basin uniquely preserve one of the earliest and most sustained records of coevolving extension and volcanism in northeastern Nevada and are the only areas where significant Oligocene extension is reported to affect upper crustal rocks (Figs. 8, 9; Rahl et al., 2002). However, interpretations of the origin of Copper Basin continue to evolve. The Copper Mountains expose a Late Cretaceous greenschist to amphibolite facies terrain intruded by a large, ca. 110 Ma quartz monzonitic stock and bounded to the east by two normal faults that juxtapose the footwall against a 1.5-km-thick sequence of volcanic and sedimentary rock in adjacent Copper Basin. 40Ar/39Ar dates on volcanic rocks from the Eocene Dead Horse Formation range from 46.3 ± 1 Ma to 37.4 ± 0.23 Ma near the base and top of the sequence, respectively (Table 1). Locally, the Dead Horse Formation includes exposures of channel-filling conglomerate, and the similarity in tephro-stratigraphic sequence with paleovalley fills farther south and west led Henry (2008) to hypothesize that Copper Basin initiated as a paleovalley that was subsequently dammed due to the onset of late Eocene extension. Therefore, lacustrine deposition in the late Eocene may record the onset of basin subsidence associated with initiation of the Copper Creek normal fault (Figs. 9A–9C). Despite the evidence for an Eocene onset of extension, footwall-derived detritus did not appear until deposition of the overlying Oligocene Meadow Fork Formation (Fig. 9D). Coats (1964) interpreted the Meadow Fork Formation to have

Figure 8. Retrodeformed cross-section sequence for the Copper Basin fault system (from Rahl et al., 2002). (A) Initial state immediately before the onset of extension at ca. 42 Ma. Due to extensive cover and poorly known or ambiguous contact relationships, structural levels above the Tennessee Mountain Formation are illustrated schematically. However, unit thicknesses and structural relationships are consistent with relationships in adjoining areas (e.g., Coats, 1964, 1987; Bushnell, 1967; Coash, 1967; Ketner et al., 1993, 1995). All units above the modern erosional surface are shaded. (B) Cross section after displacement on the Copper Creek and Meadow Fork normal faults but immediately before activity on the Bruneau Valley normal fault. Levels above the restored position of the modern erosional surface are again shaded. (C) Modern cross section established after eruption of the Jarbidge Rhyolite and down-faulting of the western fault block on the Bruneau Valley normal fault.

Cenozoic extension in the northern Great Basin

Qal

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Granitoid (Cretaceous) Intrusion (Jurassic)

TRs PPs PMh DOs OCt

Sedimentary rocks (Triassic) Sedimentary rocks (Permian-Pennsylvanian) Sedimentary rocks (Pennsylvanian-Mississippian) Sedimentary rocks (Devonian-Ordovician) Sedimentary rocks (Ordovician-Cambrian)

CZqp Quartzite and Phyllite (Cambrian-Proterozoic)

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Meadow Fork Formation (32.5 Ma to <29.5 Ma)

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44 Henry et al.

Cenozoic extension in the northern Great Basin accumulated in a narrow, northeast-trending, fault-bounded basin because of its coarse clasts and the fact that both it and the Dead Horse Formation wedge out abruptly. This formation contains rare, probably distally sourced pyroclastic-fall tuffs dated at 32.5 ± 0.2 Ma and 29.4 ± 0.4 Ma near the base and in the middle of the sequence, respectively (Table 1). These dates indicate a ca. 5 Ma hiatus in basin filling and presumably in normal faulting (Fig. 9). Both formations were subsequently tilted ~25° toward 220°. The tilted strata are intruded by 16.5 Ma basalt, and the tilted strata, the basalt, and the bounding Meadow Fork fault to the west are overlain by ca. 16 Ma Jarbidge Rhyolite. Although buried by Quaternary landslide deposits, the contact between flat-lying rhyolite and tilted Eocene-Oligocene deposits and the change of underlying strata from Dead Horse Formation northward to Meadow Fork Formation leave little doubt that the contact is an angular unconformity. Moreover, the Jarbidge Rhyolite rests nonconformably on lower plate rocks of the Coffeepot stock north and west of Copper Basin, clearly implying that the rhyolite postdated the tectonic juxtaposition of hanging wall and footwall rocks. Despite the above relationships, Eocene and middle Miocene strata in the Jarbidge Wilderness area farther east are broadly concordant. Thus the angular discordance between Eocene and Middle Miocene strata appears to be localized along the Copper Creek fault system. A steep, west-dipping normal fault in the

Figure 9. (A) Annotated oblique northeast Google Earth view of Copper Basin in north central Elko County. The basin is bounded to the west by two subparallel, SE-dipping normal faults, with Neoproterozoic to Lower Cambrian quartzite and phyllite in the footwall of the lower, Copper Creek fault, and inferred Middle Cambrian to Ordovician marble and calc-silicate rock (Tennessee Mountain Formation, Otm) in the aureole of a Late Cretaceous pluton between the Copper Creek and Meadow Fork faults. Overlying both faults is the Eocene Dead Horse Formation and conformably overlying Oligocene Meadow Fork Formation, both of which are equally tilted ~25° NW. Tuffs in the Dead Horse Formation are dated between 46.3 ± 1.0 Ma and 37.4 ± 0.2 Ma near the top of the sequence (Table 1). Lacustrine strata preserving mixed deciduous and coniferous leaf fossils lie immediately above a tuff with an age of 39.8 ± 0.2 Ma. The age of the overlying Meadow Fork Formation is constrained by ages of 32.5 ± 0.2 Ma and 29.3 ± 0.4 Ma on pyroclastic-fall tuffs near the base and middle of the sequence, respectively. The sequence is intruded by hypabyssal, plagioclase-phyric basalt with an age of 16.5 ± 0.2 Ma, and relatively flat-lying Jarbidge Rhyolite (ca. 15.5 Ma) appears to overlie the entire assemblage with angular unconformity. (B) View NNE along strike of the Copper Creek normal fault system. (C) View north of white Eocene pyroclastic-fall deposits (Dead Horse Formation, Tdh) forming NW-dipping beds below gently dipping middle Miocene Jarbidge Rhyolite (Tjr) in Copper Basin. Note the apparent angular unconformity between the Eocene and Miocene units, although the contact is covered by Quaternary landslide and colluvial deposit. (Photo by A.J. McGrew.) (D) Typical outcrop character of Meadow Fork conglomerate. The base of the Meadow Fork Formation is marked by the first appearance of abundant footwall clasts, including quartzite, calc-silicate rock, and granitic clasts that strongly resemble the Cretaceous Coffeepot quartz monzonite.

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Bruneau Valley west of Copper Mountain cuts and rotates the Jarbidge Rhyolite and other Miocene rocks and must be Miocene in age. Taken together, these relationships imply an extensional history resembling that illustrated in the retrodeformed cross sections shown in Figure 8. 40 Ar/39Ar K-feldspar and (U-Th)/He apatite ages from lower plate rocks further constrain the progression of extensional unroofing outlined above (McGrew et al., 2007). A boulder of quartz monzonite collected from the Meadow Fork Formation documents that the footwall clast cooled through (U-Th)/He apatite closure (nominally 70 °C) by 43.0 ± 2.1 Ma followed by exhumation to the surface, erosion and redeposition in the adjoining basin before 29.5 Ma. Similarly, granitoid samples from the structurally highest, western part of the lower plate yield (U-Th)/ He apatite cooling ages of 40.9–42.1 Ma, consistent with Eocene unroofing of these rocks as illustrated in Figure 8A. In addition, the structurally deeper eastern part of the range yields (U-Th)/He zircon ages of 46.1 and 53.7 Ma, and K-feldspar multi-diffusion domain modeling suggests that final cooling began in the same 40–45 Ma time frame. Finally, granitic rocks from the intermediate fault slice between the Copper Creek and Meadow Fork faults yield a (U-Th)/He apatite cooling age of 32.6 ± 2 Ma. This suggests that cooling and exhumation of the fault slice immediately beneath the Meadow Fork fault probably did not occur until the Oligocene and was approximately synchronous with the second pulse of extension recorded by the deposition of the Meadow Fork Formation. A potentially surprising aspect of the relationships at Copper Basin is the absence of a growth fault relationship, i.e., fanning of dips forming angular unconformities between the Eocene and Oligocene strata. However, the absence of such a relationship does not necessarily preclude syntectonic deposition. Rather, it could merely indicate that during the initial phases of extension the proximal part of the hanging wall immediately adjacent to the fault subsided with little or no rotation and hence without developing a growth faulting relationship. Instead, most of the rotation must have occurred during the closing phases of slip on the bounding normal fault. This progression would require that the initial fault geometry was relatively planar down to mid-crustal depths (6–10 km), at which depth the fault must have become increasingly listric. Under these conditions, rotation of the proximal part of the hanging wall would not be expected to occur until it was translated down against the lower-angle part of the listric fault surface. Taken together, currently available evidence from Copper Basin indicates that extension on the adjacent fault systems must have initiated with an early pulse of normal faulting in the late Eocene, followed by a lull during the early Oligocene, and then a pulse of extension in the late Oligocene. Extension on the Copper Creek fault system must have been largely completed before the Miocene based on: (1) (U-Th)/He apatite cooling ages no younger than 27.0 Ma in the lower plate from the footwall of the fault, (2) appearance of footwall clasts in syntectonic fanglomerate deposits by 32.5 Ma, and (3) a probable angular unconformity

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between west-tilted Oligocene rocks and overlying middle Miocene Jarbidge Rhyolite. Nevertheless, more widely distributed normal faults with substantial displacement (probably ≥1 km on the Bruneau Valley fault west of Copper Mountain) cut, offset and rotate the Jarbidge Rhyolite, documenting significant middle Miocene or later extension. Evidence from Miocene Basins and Thermochronology (J.P. Colgan) Based on a synthesis of new and existing data covering an area from the East Range to the Ruby Mountains, Colgan and Henry (2009) advanced the hypothesis that—rather than a quasicontinuous process from the Eocene to the present—Basin and Range extension in north-central Nevada took place primarily during a relatively brief interval in the Miocene that began ca. 17–16 Ma and was over by 12–10 Ma, following a period of quiescence from ca. 35–17 Ma during which little or no deformation, deposition, and volcanism took place. Primary lines of evidence for this interpretation include: (1) the lack of Tertiary rocks dating from ca. 35 Ma to 17 Ma, which indicates a period of little or no surficial deformation and volcanism in the Oligocene and early Miocene; (2) the age, composition, and distribution of Miocene sedimentary rocks (Humboldt Formation and age-equivalent units), which were deposited over a very wide area beginning at essentially the same time (17– 16 Ma) and consist of coarse material shed from rising, faultbounded blocks into nearby basins; and (3) apatite fission-track and (U-Th)/He data from exhumed footwall blocks in several ranges, which document rapid cooling beginning ca. 17–15 Ma. Rapid extension during this specific time window has been documented across nearly the entire modern extent of the northern Basin and Range province and is thus a province-wide phenomenon not limited to northeastern Nevada. In northeastern and north-central Nevada, normal faults and sedimentary basins formed during the middle Miocene are crosscut by distinctly younger normal faults that are high-angle, much more widely spaced, and represent very minor horizontal extension. The major east-dipping normal faults that bound the east sides of the Ruby Mountains and East Humboldt Ranges (Day 1) are examples of this generation of structures. The onset of this faulting is younger than 10 Ma but otherwise poorly constrained. These high-angle faults have Quaternary slip and one, the “Clover Hill fault,” observable along our route to Stop 1-4, slipped on 21 February 2008, yielding a Mw 6.0 earthquake, which did not break the surface but did significant damage to the town of Wells (dePolo et al., 2011). Summary of Observations: Cenozoic Extension All authors agree that: (1) The record of deep-crustal cooling and decompression between the Late Cretaceous and early Miocene implies substantial, but incomplete unroofing of deep-crustal rocks exposed in the RMEH, but whether this unroofing was tied to simultaneous major extension at shallower crustal levels is uncertain.

(2) Extension occurred in the Eocene, but its distribution, amount, and style are uncertain. Eocene extension may have been widespread over northeastern Nevada or more narrowly focused on a few well-developed systems such as the RMEH. This episode produced small basins, thin sedimentary deposits, and mostly small stratal tilts. Whether these differences compared to Miocene extension reflect different magnitude or different style of extension is particularly uncertain. (3) The record of Oligocene volcanism, sedimentation, and surficial extension is minimally preserved at best and may be restricted to just a few isolated localities such as Copper Basin. Whether this implies that Oligocene extension was largely absent, restricted to a particularly narrow region, or proceeded under a fundamentally different tectonic regime that may not have left a clear surficial record is unknown. (4) Extension beginning in the middle Miocene was widespread throughout northeastern Nevada, and indeed, across the entire northern Basin and Range province. The magnitude of extension varied across the region, but, where large, this episode generally produced large basins, thick sedimentary deposits, and large stratal tilts. FIELD TRIP GUIDE Day 1. Northern East Humboldt Range and the Secret Creek Gorge Area, Northern Ruby Mountains, Nevada Overview of the Ruby Mountains–East Humboldt Range (RMEH) Core Complex The first part of this field excursion focuses on field relationships at a variety of structural levels in the core complex as exposed at its northern end in the East Humboldt Range (Fig. 10). The Tertiary extensional architecture of the RMEH consists of two structural tiers (Sullivan and Snoke, 2007). The upper tier consists of non-metamorphosed to weakly metamorphosed, brittlely attenuated stratified rocks. These rocks range in age from late Paleozoic to Miocene and lie above the frictional/brittle, normal-sense, Ruby–East Humboldt detachment fault. The uppermost part of the lower structural tier is a km-thick, west-rooted, Tertiary mylonitic shear zone that passes down into a “metamorphic infrastructure” (Armstrong and Hansen, 1966). The evolution of the Ruby–East Humboldt detachment fault system and its relationship to the mid-Tertiary mylonitic shear zone of the lower structural tier have not been completely deciphered despite many detailed studies of this plastic-to-brittle fault/shear-zone system (Hacker et al., 1990; Hurlow et al., 1991; Mueller and Snoke, 1993b; MacCready, 1996; McGrew and Casey, 1993; Colgan et al., 2010; Haines and van der Pluijm, 2010). Virtually all of the brittle deformation associated with the exposed fault system overprints and therefore postdates the mid-Tertiary mylonitic shear zone. Thermobarometric data from mylonitic pelitic schists indicate a P-T of 3.1–3.7 kbar and 580–620 °C (Hurlow et al., 1991)—physical conditions well beyond the onset of quartz and feldspar plasticity (Scholz, 1990).

Cenozoic extension in the northern Great Basin Thermochronological data indicate that mylonites were significantly cooler than 300 °C by the early Miocene (Dallmeyer et al., 1986; Dokka et al., 1986; McGrew and Snee, 1994), and middle to upper Miocene rocks locally occur in the hanging wall of the detachment system (Snoke et al., 1984; Mueller and Snoke, 1993b; Fig. 10). Finally, although the bulk of the thermochronologic and radiometric data suggest that mylonitization occurred in the time interval of ca. 29–21.5 Ma (Wright and Snoke, 1993), some thermochronological data and field relationships suggest that Eocene mylonitization may be related to an earlier movement history along the shear zone that was strongly overprinted by late Oligocene mylonitization (Mueller and Snoke, 1993b; Wright and Snoke, 1993; McGrew and Snee, 1994; Snoke et al., 1997; Howard, 2003). Stops 1-1 through 1-3 are along or accessed from Nevada Highway 231 (Angel Lake Highway). Take the west Wells exit from I-80, turn south onto Humboldt Avenue, then west on the Angel Lake Highway. Drive ~4.6 mi to Stop 1-1. Note: All locations are in UTM, NAD27. Stop 1-1. Overview of the Northern East Humboldt Range and Well-Exposed, Neogene Low-Angle Normal Fault in Roadcut (Figs. 10, 11) (666240 4547445, Welcome 7.5′ quadrangle) The following description is modified from Stop 7 of Mueller and Snoke (1993b) and Stop 2-1 of Snoke et al. (1997). Pull off road to the right into large parking space. From this locality, there is an excellent view westward of the high country of the northern East Humboldt Range as well as the tree-covered forerange and low-lying foothills. The high-country is underlain by amphibolite-facies metamorphic rocks and is part of the high-grade footwall of the East Humboldt metamorphic core complex (McGrew, 1992; McGrew et al., 2000). The forerange is underlain by unmetamorphosed upper Paleozoic sedimentary rocks, and the low foothills expose Tertiary sedimentary and volcanic rocks (Fig. 10). These Paleozoic and Tertiary rocks constitute part of the hanging-wall sequence of the East Humboldt Range core complex. The East Humboldt plastic-to-brittle shear zone/detachment fault system originally separated these disparate crustal levels before being offset by late, high-angle normal faults (Mueller and Snoke, 1993a). A principal feature of the metamorphic terrane is the basement-cored, southward-closing, Winchell Lake fold-nappe (Lush et al., 1988; McGrew, 1992; Fig. 5A). Chimney Rock, a prominent isolated peak (at about azimuth 235°), is composed of orthogneiss and paragneiss that form the core of this largescale recumbent fold. The tree-covered forerange consists chiefly of Pennsylvanian and Permian carbonate rocks (Ely Limestone, Pequop Formation, Murdock Mountain Formation), but also includes Pennsylvanian and Mississippian Diamond Peak Formation. A prominent high-angle normal fault separates the Paleozoic rocks from the Tertiary strata. Reddish-brown–weathering siliceous breccia (so-called jasperoid) crops out locally along the

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contact (e.g., at ~240°). Prominent outcrops in the low-lying Tertiary terrane are lens-like masses of megabreccia derived from middle Paleozoic carbonate rocks (Stop 1-2). Finally, rhyolite porphyry of the Willow Creek complex (Jarbidge Rhyolite type) forms an isolated body along the range front at about azimuth 205°, and a prominent mass of rhyolite is at ~300°. A low-angle normal fault exposed in the west roadcut (Fig. 5D) immediately north of the pull-out separates steeply dipping Miocene fanglomerate of the hanging wall from a footwall of platy marble representing Ordovician and/or Cambrian strata intruded by pegmatitic leucogranite. Red gouge separates the two units, and the top of the marble is brecciated. As the fault is traced northward along strike, a fault-bounded slice of metadolomite (Silurian and/or Devonian protolith) is present along this tectonic boundary (Fig. 10). Optional traverse. If the high country is too snow-covered for direct field examination or not accessible, metamorphic rocks are exposed between our parking spot at Stop 1-1 and the ridgeline of Clover Hill. The flank of Clover Hill to the ridgeline is a dip slope consisting chiefly of flaggy, mylonitic quartzite (CZqs) of the footwall. Locally, scattered remnants of mylonitic, impure calcite marble structurally overlie the mylonitic quartzite. Eventually, we will reach kyanite-bearing pelitic schist (Neoproterozoic McCoy Creek Group—probably McCoy Creek G unit of Misch and Hazzard [1962]), which is characterized by a GRAIL metamorphic mineral assemblage: muscovite + biotite + plagioclase + garnet + kyanite + sillimanite + rutile + ilmenite (Hodges et al., 1992; Snoke, 1992). Based on textural relationships, sillimanite (i.e., fibrolite) is a late, minor mineral phase in the metapelitic schists exposed on Clover Hill. Compositionally banded granitic orthogneiss (inferred to be equivalent to the orthogneiss of Angel Lake; Stop 1-3) and associated paragneisses occur in a complexly folded and plastically faulted zone that forms the structurally deepest level exposed on Clover Hill below the ridgeline. Return to vehicles and continue westward along the Angel Lake Highway for ~1.8 mi. Stop 1-2. Megabreccia Deposits and Surrounding Tertiary Strata (665075 4544850, Welcome 7.5′ quadrangle) Please watch out for rattlesnakes, especially around the megabreccia outcrop. The following description is modified from Stop 8 of Mueller and Snoke (1993b) and Stop 2-2 of Snoke et al. (1997). In the northern East Humboldt Range, the Humboldt Formation immediately above the middle to late Eocene volcanic and sedimentary rocks consists chiefly of thick-bedded conglomerate and sedimentary breccia (Fig. 11). These coarsegrained deposits disconformably overlie the Eocene rocks, and pebbles of volcanic rock derived from the older Tertiary deposits are a conspicuous component of the detritus near the base of the Humboldt Formation. However, this aspect quickly changes up section, and the Humboldt Formation typically consists nearly

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Figure 10. Geologic map of parts of the Welcome and Humboldt Peak 7.5′ quadrangles, East Humboldt Range–Clover Hill area, Elko County, Nevada (A.J. McGrew and A.W. Snoke, unpublished data). (Continued on following page.)

Cenozoic extension in the northern Great Basin

Figure 10 (continued).

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exclusively of coarse detritus from Paleozoic formations, chiefly the Diamond Peak, Chainman, and Guilmette formations. Megabreccia derived from the Guilmette Formation forms discontinuous lenses in the Humboldt Formation; fine-grained lacustrine limestone, locally cherty, and minor sandstone is intercalated with these coarse deposits. The roadcut immediately across from the parking place and those to the west provide a useful starting place to examine the megabreccia deposits and encasing beds that are all within what is here termed the lower Humboldt Formation (Thl) (Fig. 11). After briefly examining the roadcut exposing lacustrine limestone and conglomerate, walk west up the road (stratigraphically downward) toward the dark gray roadcuts of megabreccia derived

from the Upper Devonian Guilmette Formation. This megabreccia is one of a group of lens-like masses within the Tertiary lower Humboldt Formation, and it both overlies and underlies lacustrine limestone and conglomerate. Therefore, this megabreccia lens, as well as the others, is encased in Tertiary deposits. After examining these roadcut exposures, walk north along the megabreccia lens examining the brecciated internal structure of the deposit. Despite pervasive brecciation, the megabreccia deposits locally display vestiges of relict bedding. If time is available, walk northeastward to another megabreccia lens composed of metadolomite and quartzite derived from Devonian and/or Silurian protoliths. Similar metadolomite and quartzite are exposed as bedrock in place on Clover Hill to the east.

Northeastern East Humboldt Range Stratigraphic Section

Fluvial-lacustrine sequence

unconformity

3000

Flat-lying, post-detachment boulder conglomerate including cobbles of Willow Creek rhyolite complex and scarce vesicular basalt.

Ash-fall tuff and reworked sandstone and siltstone, pebbly conglomerate, platy lacustrine siltstone and minor algal limestone. Lower part of section records initial unroofing of non-mylonitic metamorphic complex.

Upper Humboldt Formation (Miocene)

Willow Creek rhyolite complex

_ Extrusive rhyolite lava flow (”purple rhyolite”--13.4 ± 0.5 Ma, K-Ar sanidine) which overlies volcaniclastic talus breccia shed off domiform body of 13.8 Ma rhyolite porphyry.

Fluvial sequence

2000

Extrusive rhyolitic lava flows (14.8 ± 0.5 Ma, K-Ar sanidine) and scarce silicified siltstone. Chiefly intrusive quartz-feldspar rhyolite porphyry (13.8 ± 0.5Ma., K-Ar sanidine).

Volcaniclastic siltstone, sandstone, and conglomerate, scarce limestone (commonly fossiliferous).

Thl2

Alluvial fan sequence

disconformity 1000

Conglomerate and sandstone (red to blue-grey) with local limestone intercalations.

Conglomerate and sedimentary breccia with lenses of megabreccia; local limestone intercalations.

Lower Humboldt Formation (Miocene) Thl1

disconformity

00

Meters

faulted lower contact

Volcaniclastic conglomerate and sandstone, andesitic flows and flow breccias, subordinate rhyolite.

Tvs

Middle to Late Eocene

Figure 11. Generalized stratigraphic column of Tertiary rocks in the northeastern East Humboldt Range (modified from Mueller and Snoke, 1993b).

Cenozoic extension in the northern Great Basin Continue westward on the Angel Lake Highway for ~5.3 mi to Angel Lake, where we hike a steep, 2.0 km traverse west of the lake with an elevation gain of 150 m (~500 ft) to ~2700 m (8860 ft). Stop 1-3. Angel Lake (661050 4543196, Welcome 7.5′ quadrangle) Angel Lake cirque affords an opportunity to appraise the style and conditions of metamorphism and intrusive relationships as well as the deformational character of the high-grade infrastructure of the East Humboldt metamorphic core complex. The large-scale structural architecture of Angel Lake cirque is controlled by the Winchell Lake fold-nappe (Figs. 5A, 10, 12). Although the best exposures of the closure of this fold occur at Winchell Lake cirque ~7 km to the south, a transect from Angel Lake westward to the crest of Greys Peak would take the climber completely through both limbs of the fold. The Greys Peak fold, a map-scale parasitic fold on the upper limb of the Winchell Lake fold-nappe, is visible on the eastward-facing cliff face directly beneath Greys Peak. Well-developed mylonitic fabrics overprint the fold-nappe in the upper part of the cirque, forming a thick zone of protomylonitic gneiss with WNW-directed kinematic indicators. These rocks represent the northern extent of the Ruby–East Humboldt mylonitic shear zone, which extends over 100 km to the south along the west flanks of both the East Humboldt Range and Ruby Mountains. The mylonitic rocks gradually give way to coarse-grained gneisses with increasing structural depth, and our transect is located entirely within the coarse-grained gneiss domain beneath the mylonitic shear zone. Shear-sense indicators, to the extent that they are observed in this zone, tend to be ESE-directed, antithetical to the overlying mylonitic zone (McGrew, 1992). The transect climbs through the characteristic stratigraphic sequence on the lower limb of the fold-nappe from bottom to top (Fig. 12): (1) a quartzite and schist sequence of probable Early Cambrian to Neoproterozoic age (CZqs); (2) a thin sequence of calcite marble and calc-silicate gneiss that is probably Upper Cambrian and Ordovician (OCm); (3) a discontinuous, thin orthoquartzite layer (<2 m thick) inferred to correlate with the Ordovician Eureka Quartzite (Oq); (4) a sequence of dolomite marble correlated with Ordovician to Devonian dolomite (DOd); (5) more calcite marble, locally including isolated enclaves of intensely migmatized rusty-weathering graphitic schist (MD?mu); (6) a sequence of highly migmatized quartzofeldspathic to quartzitic paragneiss and schist (Zqs) inferred to correlate with Neoproterozoic (McCoy Creek Group) based on U-Pb SHRIMP analysis of detrital zircons (W.R. Premo, 2010, personal commun.); (7) a thick sequence of distinctively banded biotite monzogranitic orthogneiss of Angel Lake that is probably Archean or near-Archean (XWog; W.R. Premo, 2010, personal commun.; Lush et al., 1988; Premo et al., 2008; McGrew and Snoke, 2010; Premo et al., 2010); and (8) a sequence of strongly

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migmatitic biotite schist and immature, quartzofeldspathic paragneiss forming the core of the Winchell Lake fold-nappe in this location and probably Paleoproterozoic or latest Archean age (XWpg) based on U-Pb SHRIMP analysis of detrital and metamorphic zircons (W.R. Premo, 2010, personal commun.). The same sequence is present in reverse order on the upper limb of the fold-nappe below Greys Peak. The contact between inferred middle to upper Paleozoic marble (MD?mu) and Neoproterozoic paragneiss (Zqs), must be a pre-folding, pre-metamorphic fault. The contact between Neoproterozoic McCoy Creek Group (Zqs) and Archean orthogneiss (XWog) could be either another pre-metamorphic fault or an unconformity. The contact between Archean orthogneiss and paragneiss (XWpg) could be unconformable, intrusive, or tectonic. Although the oldest rocks occupy the core of the Winchell Lake fold-nappe, the stratigraphic facing direction of the miogeoclinal sequence is inward, not outward. Therefore it is unclear whether the Winchell Lake fold is anticlinal or synclinal. If synclinal, the Precambrian gneiss complex must have been inverted as it was thrust over the miogeoclinal rocks before folding. If anticlinal, before folding, the miogeoclinal rocks must have been inverted and faulted over the Neoproterozoic McCoy Creek Group, which itself overlays the Archean–Early Paleoproterozoic gneiss. Either reconstruction requires a complex, polydeformational history, although the synclinal interpretation may be easier to envision. Nevertheless, the slightly greater attenuation of the lower limb of the fold and the presence of both south- and northvergent small-scale folds on the lower limb would seem to argue for an anticline as discussed below. Small-scale folds occur throughout the transect, with hinge lines parallel to well-developed stretching lineations, showing an average orientation of 5°/295°. Similarly, the axial surfaces of small-scale folds are approximately parallel to foliation, with an average orientation of 270°/15°. Three-dimensional constraints indicate that the map-scale folds (i.e., the Greys Peak fold and Winchell Lake fold-nappe) show the same WNW trend as the smaller scale structures. However, there is considerable dispersion in the small-scale structural data at deep structural levels, and a cryptic, northerly trending lineation of uncertain age can be observed locally. In addition, whereas folds on the upper limb of the Winchell Lake fold-nappe verge systematically southward toward the hinge zone of the first-order structure, folds on the lower limb verge both northward and southward and commonly exhibit “refolded fold” geometries (Ramsey Type 3 interference patterns) (Fig. 5C). The simplest interpretation of these relationships would entail the overprinting of S-folds by Z-folds as they migrate from the upright to the overturned limb of the largerscale structure as the fold-nappe grows in “tractor tread” fashion. If so, then the fold-nappe would be required to be anticlinal. The most characteristic pelitic mineral assemblage at all structural levels in Angel Lake cirque is biotite + sillimanite + garnet + quartz + plagioclase ± K-feldspar ± chlorite ± muscovite ± rutile ± ilmenite. However, scarce relict subassemblages of kyanite + staurolite survive on the upper limb of the Winchell

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Figure 12. Detailed geology of the Angel Lake area showing localities A–F (Stops A–F on map) for Stop 1-3 (modified from McGrew, 1992).

Cenozoic extension in the northern Great Basin Lake fold-nappe. In addition, late stage muscovite and chlorite become increasingly prominent in the well-developed mylonitic rocks at higher structural levels. Five samples of schist from Angel Lake cirque yield internally consistent P-T estimates ranging from 630° ± 105 °C, 5.2 ± 1.2 kbar near the base of the cirque to 700° ± 125 °C, 8.3 ± 1.2 kbar on the ridge line in the northwest corner of the cirque (McGrew et al., 2000). Mark Peters analyzed an amphibolite collected at 2780 m (9120 ft) on the north side of the cirque that yielded the highest P-T estimate reported from the East Humboldt Range, 752° ± 138 °C, 9.3 ± 1.1 kbar. Thus, the Angel Lake area by itself yields much the same P-T trend as has been recognized for the whole East Humboldt Range, and we interpret this trend to represent re-equilibration at various stages along a decompressional P-T trajectory (Fig. 6; McGrew et al., 2000). The higher structural levels in Angel Lake cirque may have cooled and equilibrated at a relatively early stage in the unroofing history as indicated by an 40Ar/39Ar hornblende cooling age isochron of 51 ± 2 Ma at an elevation of 2890 m (9480 ft) on the north side of the cirque. On the other hand, samples from deeper structural levels may have experienced a slower cooling history that resulted in equilibration at a later stage because 40Ar/39Ar hornblende cooling ages are between ca. 30 and ca. 40 Ma at elevations beneath ~2775 m (9100 ft) (McGrew and Snee, 1994). Cooling of the northern East Humboldt Range through 40 Ar/39Ar muscovite and biotite closure temperatures (nominally 350 °C and 280 °C, respectively) occurred by ca. 21.5 Ma, presumably due to exhumation along the RMEH mylonitic shear zone (McGrew and Snee, 1994). Localities A–F below (Fig. 12) mark our planned transect through the lower part of Angel Lake cirque. As snow often lingers deep in the cirque until well into June, we may have to modify this plan depending on conditions. Locality A. Follow the well-developed path around the south side of the lake to the low bedrock hill labeled “A” on the geologic map (Fig. 12). This locality exposes the inferred Neoproterozoic and Cambrian quartzite and schist (Zqs) at the deepest structural levels on the lower limb of the fold-nappe. The East Humboldt Range is permeated by intrusive rocks of various ages and compositions, particularly abundant leucogranite at this deep structural level. We interpret at least four generations of leucogranitic intrusion based on cross-cutting relationships. In addition to the leucogranites, sheets of biotite monzogranitic orthogneiss are abundant throughout the RMEH and have yielded U-Pb zircon ages of ca. 29 Ma in many locations, including one at the northwest corner of Angel Lake (locality F) (Wright and Snoke, 1993). Consequently, these monzogranitic sheets are useful gauges for deciphering the structural chronology. The biotite monzogranites cut older generations of leucogranite, but some leucogranite either cuts or intermingles with the monzogranites. At higher structural levels, mylonitic fabrics clearly overprint the monzogranites, providing a crucial constraint on the age of mylonitic deformation. Steeply dipping, amygdaloidal basalt dikes that occupy three deep clefts through the south wall of the cirque were the final phase in the intrusive history of

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the East Humboldt Range. The end of one of these amygdaloidal basalt dikes can be inspected near this locality. Similar dikes yield approximate ages of 17–14 Ma elsewhere in the Ruby Mountains and East Humboldt Range, providing a younger limit on the age of plastic deformation (Snoke, 1980; Hudec, 1992). Locality B. Continue over and around the small hill on the southwest side of the lake and cross the stream to point “B” on the map (Fig. 12). The rocks near the base of the slope form a particularly diverse and structurally complex paragneiss sequence. Can you find any amphibolite boudins near here? Normally, amphibolite bodies in the East Humboldt Range are found only in the older Precambrian rocks forming the central part of the cirque, where they probably represent metamorphosed mafic bodies that were presumably intruded before deposition of the upper part of the McCoy Creek Group and the overlying Prospect Mountain Quartzite. However, here one or two amphibolite bodies occur in the inferred Cambrian and/or Neoproterozoic quartzite and schist sequence. The age of the strata at this locality is constrained by anomalously high δ13C values in the thin marble layers that are interdigitated with the quartzite and schist. As discussed in detail by Peters and Wickham (1992) and Wickham and Peters (1992), the high δ13C values of these marble layers probably are original protolith values and indicate deposition during one of the Neoproterozoic carbon isotope excursions (e.g., Halverson et al., 2005). Accordingly, these rocks are inferred to correlate with the Neoproterozoic McCoy Creek Group. Locality C. Contour along the foot of the slope to the base of the waterfall on the west side of the cirque; cross to the south side of the stream. The traverse works upward from the inferred Prospect Mountain Quartzite at the base through the lower calcite marble sequence (OCm) to a distinctive thin, white orthoquartzite approximately halfway up the waterfall. This is the Ordovician Eureka quartzite (Oq), and immediately above it is a relatively thick unit of dolomitic marble (DOd). Marble and calc-silicate assemblages here typically consist of calcite + diopside + quartz ± dolomite ± phlogopite ± plagioclase ± grossular ± scapolite ± K-feldspar ± sphene ± amphibole ± epidote. Peters and Wickham (1994) report that the diopside-bearing primary assemblages probably equilibrated at ≥6 kbar, 550–750 °C and likely record conditions during Late Cretaceous or early Cenozoic metamorphism. They also report a secondary subassemblage, amphibole + grossular + epidote, that records infiltration of water-rich fluids under a metamorphic regime that proceeded from high temperature (600–750 °C) to lower temperature (<525 °C) conditions. This event was probably related to Tertiary extension and associated magmatism. Above the dolomitic marble is more calcite marble, presumably of Devonian to Lower Mississippian age (MD?mu). At the top of the waterfall are a series of steep, ~1-m-thick, cross-cutting aplitic leucogranite dikes, probably the latest intrusive phase exposed in the East Humboldt Range other than the middle Miocene basaltic dikes. Locality D. From the top of the waterfall, regroup at a ledge on the north side of the stream marking the contact with the overlying quartzite and schist sequence (Zqs). Detrital

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zircon data show a well-developed spike of Grenville-aged zircons, with the youngest zircons being ca. 900 Ma (W.R. Premo, unpublished data). Accordingly, this paragneiss probably correlates with the Neoproterozoic McCoy Creek Group but contrasts with the quartzite and schist sequence at the base of the cirque in that it is dominated by impure, feldspathic and/or micaceous metapsammite with little “clean” quartzite and almost no marble, except for probable infolds of middle to upper Paleozoic marble directly adjacent to the contact. In addition, volumetrically small but relatively common and widespread, inferred originally mafic intrusions now form sheets and boudins of orthoamphibolite. Since these orthoamphibolite bodies are absent from the inferred Paleozoic metasedimentary rocks and rare in the inferred upper McCoy Creek Group rocks such as those described at Stop B, we infer that these rocks represent an older part of the McCoy Creek Group that may not be exposed elsewhere in the Great Basin. The contact here between the McCoy Creek Group paragneiss and the underlying, inferred middle to upper Paleozoic marble sequence must be a pre-metamorphic, pre–Winchell Lake fold tectonic contact, and may be one of the major structures that buried the Paleozoic rocks to such great depth. Could the rocks at the contact be annealed mylonite? Can you recognize any possible vestiges of this inferred pre–Late Cretaceous shearing event? Locality E. The traverse from locality D works through sporadic outcrops of lower McCoy Creek paragneiss to locality E on the north side of the cirque, where the contact with the banded biotite monzogranitic orthogneiss of Angel Lake (XWog) is exposed (Fig. 5B). Generally pale gray, this coarse-grained rock locally displays feldspar augen and is easily recognized by its distinctive striped appearance due to biotite segregation. As discussed in detail in the section on Precambrian, this unit was dated as Late Archean by Lush et al. (1988), but U-Pb SHRIMP analysis of zircons from a newly collected, more homogeneous and far less migmatized sample near the top of the orthogneiss complex west of Chimney Rock yielded an earliest Paleoproterozoic age (2450 ± 5 Ma; W.R. Premo, 2010, personal commun.). In addition, renewed field investigation of the gneiss complex of Angel Lake led to the delineation of a previously unmapped unit of impure, micaceous, feldspathic paragneiss and intensely migmatized biotite schist occupying the core of the Winchell Lake fold (XWpg; Fig. 12). Newly acquired U-Pb SHRIMP data on detrital zircons from this unit constrain it to be younger than ca. 2550 Ma (the age of the youngest detrital zircons) and older than a previously unrecognized ca. 1.8 Ga metamorphic event recorded by a few metamorphic zircons (W.R. Premo, 2010, personal commun.). So far we have not been able to decipher the original contact relationships between this newly recognized late Archean or Paleoproterozoic paragneiss and the orthogneiss that is folded around it, but if time and snow conditions permit, ambitious climbers may be able to climb to inspect this newly recognized map unit for themselves. Many important questions remain to be resolved about the only Archean or near-Archean rocks exposed in Nevada, and the collaboration with Wayne Premo is continuing to improve understanding of the history of these complicated exposures.

Locality F. From locality E we will work our way down the north side of the cirque to locality F at the northwest corner of Angel Lake. As you walk, try to decipher the relative age relationships between leucogranitic intrusions, and note the variation in leucogranite abundance depending on the host rock type. Some leucogranites are fully involved in folding whereas others cut folds. Outcrops at locality F consist predominantly of biotite monzogranitic and leucogranitic orthogneiss. The biotite monzogranite at this locality has yielded a U-Pb zircon age of 29 ± 0.5 Ma (Wright and Snoke, 1993). Is this sheet of monzogranite folded? Walk around the corner of the outcrop before you decide. Some monzogranitic sheets are clearly involved in folding, but others cut folds, and at map scale a number of monzogranitic bodies cut the Winchell Lake fold-nappe itself, lending credence to the interpretation that the Winchell Lake fold predates Oligocene deformation. However, the monzogranitic orthogneisses bear the same WNW-trending stretching lineations as the country rock, and at higher structural levels they are overprinted by mylonitic microstructures, thus documenting a post–29 Ma age for extensional deformation. It seems highly likely that older structures were profoundly transposed during Tertiary deformation, including the Winchell Lake fold-nappe itself. From locality F, follow the trail around the north and east sides of Angel Lake to the parking lot. Return to Wells and drive ~28 mi south on U.S. Highway 93 from the east Wells exit off I-80. Turn west on Nevada Highway 229, which turns north after ~14 mi. Drive ~1 mi north to Stop 1-4. Stop 1-4. Secret Creek Gorge, Northern Ruby Mountains (647160 4525050, Soldier Peak 7.5′ quadrangle) Pull left into large parking space. From here we walk back up the road to Stop 1-4. The following description is modified from Stop 12 of Snoke et al. (1984) and Stop 3-1 of Snoke et al. (1997). Please watch out for rattlesnakes in this area. This stop consists of a guided traverse through an anastomosing system of distinctive lithologic slices bounded by lowangle normal faults (perhaps best referred to as an extensional duplex structure; Fig. 13). The purpose of this traverse is to demonstrate the complex structural style characteristic of the lowangle fault complex, but also to develop the structural chronology between mylonitic deformation, low-angle normal faulting, and high-angle normal faulting. To facilitate the use of this guide, specific localities are designated A–F on Figure 13. Locality A (roadcut along Nevada 229). Mylonitic, interlayered migmatitic schist and impure quartzite with subordinate orthogneiss, cut by numerous westward-dipping normal faults. Many of the normal faults are planar, but a few are clearly curviplanar. Associated with the westward-dipping normal faults are spectacular drag features as well as crushed zones and thin ultramylonitic to cataclastic layers along the fault planes. In addition, flaggy micaceous quartzites with conspicuous mica porphyroclasts (“mica fish”) are useful indicators of the senseof-shear in the mylonite zone. Other mesoscopic criteria useful

Cenozoic extension in the northern Great Basin in the determination of sense-of-shear include asymmetric feldspar porphyroclasts, composite planar surfaces in pelitic schists (S-C-C′ fabric), and mesoscopic folds that deform the mylonitic foliation. Well-developed microstructural fabrics are also common in these mylonitic rocks (e.g., the mylonitic impure quartz-

115o16’W

o

115 15’W

55

ites are classic examples of Type II S-C mylonites of Lister and Snoke, 1984). All of these criteria taken together indicate a top to the west-northwest sense-of-shear throughout the quartzite and schist unit in the northern Ruby Mountains and southwestern East Humboldt Range.

o 115 14’30”W

o 115 14’W

115o13’30”W

Figure 13. Geologic map of the Secret Creek gorge area, northern Ruby Mountains, Nevada (modified from Snoke et al., 1997). Qa—alluvium (Quaternary); Qoa—older alluvium (Quaternary); Qls—landslide deposits (Quaternary); Ts—sedimentary rocks (Miocene); Tr—rhyolite (middle Miocene); Tgn—granitic orthogneiss (Tertiary); Pp—Pequop Formation (Permian); Plc—limestone and conglomerate (Permian); Pu— Permian rocks undivided; Pe—Ely Limestone (Pennsylvanian); PMdp—Diamond Peak Formation (Pennsylvanian and Mississippian); Dg— Guilmette Formation (Devonian); DOd—metadolomite (Devonian to Ordovician); Oe—Eureka metaquartzite (Ordovician); hc—Horse Creek assemblage (metasedimentary and granitic rocks); CZqs—impure metaquartzite and schist (Cambrian and Neoproterozoic). Standard symbols for attitude of bedding or foliation, trend and plunge of lineation, hinge line of mesoscopic fold, geologic contacts, and normal faults (except double tick marks on the upper plate of brittle low-angle normal fault except double tick marks on the upper plate of brittle low-angle normal fault [detachment fault] and filled squares on upper plate of plastic-to-brittle low-angle normal fault). Faults are dotted where concealed. Localities A–F are field trip stops.

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The mylonitic quartzites at this locality as well as other rock types in the Secret Creek gorge area were the subject of a stable isotope study by Fricke et al. (1992), which demonstrated the importance of meteoric water infiltration during mylonitization. Locality B (roadcut exposure along Nevada 229, west of Locality A). We have crossed a low-angle normal fault that separates the overlying Horse Creek assemblage from the underlying quartzite and schist unit. The Horse Creek assemblage is diverse and includes impure calcite marble and calc-silicate gneiss and schist (inferred Ordovician and Cambrian protoliths), and mafic to felsic orthogneisses including a distinctive deformed biotitehornblende mafic quartz diorite. Mylonitic rocks are ubiquitous and include spectacular calc-mylonite containing competent mineral grains and rock clasts. Many of the obvious folds deform the mylonitic foliation and display westward vergence. Folded boudins and dismembered folds are other manifestations of a complex strain history (inferred as progressive, non-coaxial deformation). Scarce sheath folds also occur in the Horse Creek assemblage. Locality C (slope above old road bed). Folded mylonitic white quartzite in the Horse Creek assemblage. This west-vergent fold displays rotation of an earlier mylonitic lineation during late folding in the mylonitic shear zone. Locality D (along the old road). Disharmonic west-vergent folds in the Horse Creek assemblage. Rock types include impure calcite marble with mafic pelitic layers, granodioritic augen gneiss, and white quartzite. Note how layer thickness controlled the amplitude and wavelength of the folds. The traverse between localities D and E along the old road crosses the low-angle fault contact between the overlying Horse Creek assemblage and the underlying quartzite and schist unit several times. A secondary black ultramylonite is common along this contact (i.e., a ductile-brittle low-angle fault) where the mylonitic foliation in both the upper and lower plates is roughly subparallel. In other cases, where foliation in the upper plate is highly discordant to foliation in the lower plate, the contact appears to be a brittle low-angle normal fault that has perhaps soled into (i.e., reworked) the earlier ductile-brittle low-angle fault. Furthermore, steeper normal faults cut the Horse Creek plate. Locality E. We are presently situated on the Horse Creek assemblage. The contact between the Horse Creek assemblage and underlying quartzite and schist unit can be seen in fine detail up canyon. The mapped low-angle fault contact is probably composite; parts of the contact are a low-angle, ductile-brittle fault and parts are younger, low-angle brittle normal faults that have soled into the older ductile-brittle fault. Locality F. The slope above the old road crosses two lowangle normal faults. The lower fault (the main break in metamorphic grade and here the top of the mylonitic zone) separates the Horse Creek assemblage from a higher slice of low-grade metadolomite; the upper fault separates the metadolomite from a structurally higher slice of low-grade but fossiliferous Upper Devonian Guilmette Formation. At this locality, the Guilmette Formation is a medium gray, highly calcite-veined, low-grade metalimestone. A similar tectonic slice of Guilmette Formation exposed south of

Secret Creek gorge yielded Late Devonian conodonts with color alteration index (CAI) = 5½, suggesting that the host rock reached 300–350 °C (A. Harris, 1982, written commun.). At the top of the hill (i.e., the resistant gray exposures of Guilmette Formation), part of a roadcut that exposes dismembered Diamond Peak Formation can be seen to the east-southeast, across from where the vehicles are parked. This roadcut is topographically and structurally above your present position. Therefore, another low-angle fault must separate the Upper Devonian Guilmette Formation from the Mississippian and Pennsylvanian Diamond Peak Formation. Structurally above the roadcut, Miocene Humboldt Formation is also in low-angle fault contact with the underlying rocks of the Diamond Peak Formation as well as Horse Creek assemblage (Fig. 13). Haines and van der Pluijm (2010) used X-ray diffraction to characterize clay minerals in gouges related to low-angle and high-angle faults exposed along Nevada 229 and 40Ar/39Ar to date the clay minerals. A key locality in their study was the roadcut across from where the vehicles are parked (their locality Secret-2). Authigenic illite-rich illite/ smectite in the gouges, including this locality, yielded ages of 11.6 ± 0.1 Ma, 12.3 ± 0.1 Ma, and <13.8 ± 0.2 Ma. Based on these ages, Haines and van der Pluijm (2010) concluded the last major period of activity along the low-angle faults (part of the detachment system) and a kinematically related high-angle fault was 13–11 Ma. Upon completion of the traverse, please congregate in the parking area and examine the roadcut across from the vehicles. Return to Wells by continuing westward to the Halleck interchange and then eastward on I-80. Stop 1-4 can also be accessed by this return route. Day 2. North Paleovalley and Copper Basin In dry conditions, Copper Basin can be reached by taking Elko County Road 747 (the Charleston-Deeth Road) 42.7 mi north from I-80 at the Deeth exit (Nevada Highway 230) to the intersection with Elko County Road 746 at Charleston Reservoir (Stop 2-3 is in the low hills northwest of this junction). In poorer weather, take Nevada Highway 225 (the Mountain City Highway) ~53.6 mi north from Elko to Elko County Road 746. Take ECR 746 ~20.8 mi east to the intersection with ECR 747 at Charleston Reservoir. By either route, head north on the road to Jarbidge (Jarbidge-Charleston County Road; National Forest Development Road 062). Drive ~13.1 mi north to an unmarked track to west (626756 4622232), the start of an ~6 km traverse (round-trip; Fig. 14). Stop 2-1. Copper Basin The traverse leads westward along the south ridgeline of Copper Basin, which provides an excellent perspective of the geology of Copper Basin. Copper Basin is an ~25 km2 area of Eocene ash-flow tuff and tuffaceous, partly lacustrine sedimentary rocks (the Dead Horse Formation) and Oligocene conglomerate (the Meadow Fork Formation) in the hanging wall of the

Cenozoic extension in the northern Great Basin

57 41°47’

Tjr

OCt

Meadow Fork Creek

Tmf

OCt

54 20

Tmf

24 15 26 X 35 25 30

32.5±0.2 X31 Q X31

CZ

Q 18

r 40

16 33

31

35

Ho

47.3±0.2 Ts1 X Tsb, 16.6±0.2 X

26

15

Tsb

24

X flora Tbcc? 39.8±0.4

Tdh

35

35

CZ

?

k ee Cr er pp Co

? Tpb? 41.6±0.2 X Tt1 X

41°44’

Tpb? 43.8±0.9

Tsb

Steens-type basalt

N

Oligocene Meadow Fork Fm

PMi

+Tbcc Tt1

Q

Ma

PPPs

?

tuffs and sediments

CZ

Quartzite and phyllite

ty Se ve n

Tennessee Mtn Fm

?

OCt

Tjr 25

Pre-Cenozoic Upper Paleozoic and upper JdRoberts Mtn allochthon plate of

Fault, dashed where approximately located Strike and dip, 27 measured Strike and dip direction, interpreted from 41°43’ aerial imagery

Ts1

Tt1

Tjr 0

115°30’

Tjr Tjr?

Pz

Eocene Dead Horse Fm Tt1 Ts1 and undivided

Pz

Tjr

?

Jarbidge Rhyolite DOs

jor

Figure 14

Quaternary Miocene

Tjr

Tdh

Gu

lch

Pz

Tjr

Approximate parking and field trip traverse Q Stop 2-1

Tt1? Andesite 42.7±0.1 X Pz

Ts1

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T45, 44.90±0.07

Pz

Tmf

41°45’

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40

?

45

X Tdh

De

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Q

14

14

32

29

ad

41°46’

Tjr

Q

13

Qtz monzonite clast 37 20 20 biotite 40/39 = 84 Tmf apatite (U-Th)/He = 43±2

Q 37.6±0.4 k 50 ee X r C Tdh 41.5±0.2 se

Cr ee k

29.3±0.4

43 38 34

x

Cop per Cr

CZ

Tmf 21

36

eek fault

Copper Mountain

Q

45

70 40

Si

CZ

Sanidine 40Ar/39Ar (Henry, 2008) X Biotite 40Ar/39Ar, Apatite (U-Th)/He (Rahl et al., 2002; McGrew and Foland, unpublished data) All ages in Ma 1

Tjr 2 km

115°28’

Figure 14. Simplified geologic map of Copper Basin (from Coats, 1964, 1987; Rahl et al., 2002; and this study). T45—45 Ma tuff; Tpb—plagioclase-biotite tuff; Tbcc—tuff of Big Cottonwood Canyon. See text for discussion and Figure 2 for location.

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Copper Creek fault, a major east-dipping normal fault (Figs. 9, 14; Table 1; Coats, 1964; Axelrod, 1966a, 1966b; Rahl et al., 2002). In the northeastern part of the basin, coarsely plagioclasephyric basalt dated at 16.5 ± 0.2 Ma intrudes the Dead Horse Formation, and middle Miocene Jarbidge Rhyolite overlies the Eocene-Oligocene rocks in what appears to be an angular unconformity (Fig. 9; Coats, 1964; Rahl et al., 2002). The Dead Horse Formation is at least 800 m thick and consists of ash-flow tuff overlain by and interbedded with tuffaceous, lacustrine sedimentary rocks and lesser basal conglomerate and andesite. At least four ash-flow tuffs are exposed in the southern part of the Dead Horse Formation along the traverse: the 45 Ma tuff (625926 4622238), a plagioclase-biotite-hornblende tuff (626292 4622357), the ca. 41 Ma tuff of Coal Mine Canyon, and the 40 Ma tuff of Big Cottonwood Canyon (625304 4622726; Fig. 14; Henry, 2008). A channel along the ridgeline (625436 4622741) cuts into the tuff of Big Cottonwood Canyon and contains pebbles up to 4 cm in diameter of Paleozoic quartzite and chert and Tertiary volcanic rocks, mostly the tuff of Big Cottonwood Canyon. Paleozoic clasts are mostly well rounded, whereas Tertiary clasts are subrounded to subangular. From the channel fill proceed west along the ridge to the Copper Creek fault (625010 4622587) and follow the fault northward to Copper Creek. Neoproterozoic to Lower Cambrian Prospect Mountain Quartzite crops out in the footwall. The hanging wall of the fault consists of Dead Horse Formation on the ridgeline, but a fault slice of Meadow Fork Formation overlies the fault in the valley. The Copper Creek fault has an average attitude of 025°/35°; average slickenlineations plunge 34° toward 123°, indicating nearly ideal dip-slip. The overlying strata are oriented on average 220°/24°, with both the Dead Horse and Meadow Fork Formations showing statistically identical dips. Restoring the Cenozoic strata to horizontal rotates the Copper Creek fault to 030°/58° (Fig. 8; Rahl et al., 2002). Based on this restoration, Rahl et al. (2002) estimated 8–12 km of displacement on the Copper Creek fault system, exhuming the Copper Mountains from a paleodepth of 11 ± 3 km as the net result of Eocene and Oligocene extension. Good exposures of the contact between the Dead Horse Formation and the overlying Meadow Fork Formation are along Copper Creek at the mouth of Dead Horse Creek (625738 4623476) and of the lower Meadow Fork Formation for a few hundred meters north along Copper Creek from there. The Meadow Fork Formation is at least 600 m thick and consists of tuffaceous sedimentary rocks and coarse arkosic sandstone and conglomerate containing clasts of quartzite, marble, phyllite, and granitoids up to 1 m in diameter. Biotite 40Ar/39Ar dates on rare, interbedded pyroclastic-fall tuffs are 32.5 ± 0.2 Ma near the base and 29.3 ± 0.4 Ma in the upper middle part of the section (Fig. 14; Table 1). Climb uphill toward the east and southeast to the Copper Basin Flora locality (626590 4622994), a mixed deciduous-coniferous flora discovered by Axelrod (1966a, 1966b) and deposited in a marginal lacustrine facies of the Copper Basin (Rigby et al., 2006). Paleoelevation estimates from the flora range from 1.1 km

(Axelrod, 1966a, 1966b), to 2.0 ± 0.2 km (Wolfe et al., 1998), and to 2.8 ± 1.8 km (Chase et al., 1998). Current elevation of Copper Basin is ~2.1 km. The flora are in a thin (~1 m) horizon of light brown, organic-rich shale. Approximately 7–10 m beneath the shale are interbedded tuffaceous siltstone, sandstone (commonly in graded beds), and clast-supported conglomerate, consisting mostly of angular, 2–10-mm diameter volcanic clasts with subsidiary quartzite and phyllite in an ashy matrix. Less than 2 m below the base of the conglomerate is an ~6-m-thick, massive, white, fine-grained tuff with sparse small phenocrysts of biotite and feldspar. The tuff yields an 40Ar/39Ar age of 39.8 ± 0.2 Ma (Table 1), suggesting a possible correlation with the tuff of Big Cottonwood Canyon. If so, the massively bedded, cobble conglomerate filling a channel above that tuff is missing here, replaced by a few graded beds of tuffaceous pebble conglomerate. This location provides a good view of the northern part of the Dead Horse Formation, which consists of lacustrine deposits containing a series of white, fine-grained pyroclastic-fall deposits. Although Rahl et al. (2002) reported that these pyroclasticfall deposits overlay the ash-flow tuffs of the southern part of the basin, new 40Ar/39Ar dates indicate that the fall deposits overlap in age with the ash-flow tuffs (Fig. 14; Table 1). Moreover, the fall deposits appear to be along strike with the dated locations of the 45 Ma tuff and tuff of Big Cottonwood Canyon. A fall deposit near the base of the section yields a biotite plateau age of 47.3 ± 0.2 Ma that would make it the oldest known Eocene tuff in northeastern Nevada. An overlying tuff dated at 41.5 ± 0.2 Ma probably correlates with the plagioclase-biotite tuff, and the tuff just below the Copper Basin flora may correlate with the 40 Ma tuff of Big Cottonwood Canyon. The top of the Dead Horse Formation is marked by a pyroclastic-fall tuff dated at 37.6 ± 0.4 Ma (Rahl et al., 2002). Based on these data, A.J. McGrew interprets the fall deposits to have accumulated in a lake mostly contemporaneously with deposition of the terrestrial ash-flow tuffs. In this better studied, northern part of the basin, the Dead Horse and Meadow Fork Formations mostly strike northeast and dip moderately northwest, although attitudes vary widely, indicating an angular unconformity with overlying Jarbidge Rhyolite (Fig. 14). Based only on examination of air photos and Google Earth images, the Eocene-Oligocene(?) rocks strike north and dip moderately to the east in the southeastern part of the basin. The Jarbidge Rhyolite appears to strike and dip the same way. We infer that steep, west-dipping Miocene faults cut and tilt all the rocks in the eastern part, resulting in a northeast-striking, extensional-accommodation anticline between the two differently dipping parts of the basin. This eastern tilt domain may be visible from the flora locality or from the parking area. Climb back uphill to the vehicles and drive ~12 mi south toward Charleston Reservoir along the Jarbidge-Charleston Road. Stop 2-2 (Optional). Jarbidge Rhyolite Lava Flow Margin Please watch for rattlesnakes. Drive south ~12 mi toward Charleston Reservoir along the Jarbidge-Charleston road. The hills to the east are the 45 Ma

Cenozoic extension in the northern Great Basin tuff (0623727 4609661) and Jarbidge Rhyolite, while most of the sedimentary deposits to the west are tuffaceous sedimentary rocks (Humboldt Formation?; Coats, 1987). Approximately 1.5 mi north of the junction with Elko County Road 746 (at Charleston Reservoir), stop and park (0624540 4606261). The outcrops immediately east of the road provide an outstanding example of a rhyolite lava flow-top (carapace) breccia (Fig. 15A). Jarbidge Rhyolite lavas typically contain 15%–40% smoky quartz and feldspar phenocrysts and are meta-to-peraluminous (~72–78 wt% SiO2 and average TiO2/MgO = 8.2). These and other chemical characteristics identify Jarbidge magmas as “A-type.” The rhyolite breccia consists of abundant pebble to boulder-sized clasts of rhyolite vitrophyre, flow-banded stony rhyolite, and massive stony rhyolite. Vitrophyre boulders weather out from beneath the breccia in places; this is consistent with the breccia being the upper carapace of the flow. Callicoat (2010) interpreted this exposure to be the western margin of a Jarbidge Rhyolite lava flow. The total volume of Jarbidge Rhyolite exposures in this region is estimated to be ~500 km3, which is similar to the Central Plateau rhyolites in Yellowstone (Callicoat et al., 2010). However, unlike Yellowstone, no caldera-source appears present. Instead, Callicoat (2010) suggested that the Jarbidge magmatic system, although very large (total magma volumes equal or greater than 1500 km3), was characterized by primarily effusive eruptions similar to other large volume silicic systems dominated by effusive volcanism (e.g., the Taylor Creek Rhyolite, New Mexico; the Badlands lava flow, Idaho). Jarbidge rhyolite lava flows between this location and Copper Basin, as well as flows farther north along the road, yield ages between ca. 16 and 15.5 Ma (Callicoat, 2010; Brueseke, unpublished data). Stop 2-3. Charleston Reservoir Fluvial Megabreccia and Ash-Flow Tuff Relationships Please watch for rattlesnakes. Drive south to the Elko County Road 746–JarbidgeCharleston Road junction and park in the open area east of the junction. Sagebrush-covered flats and low hills south, north, and west of the parking area are underlain by Cenozoic sedimentary strata and ash-flow tuffs (Fig. 15B). Coats (1987) mapped the ash-flow tuffs as part of a regionally widespread Eocene unit (Tt1). The sediments are likely part of the younger Humboldt Formation. Massive outcrops adjacent to and east of the reservoir are Jarbidge Rhyolite lava flows (Callicoat, 2010). Walk across the road(s) and head northwest toward the megabreccia outcrops on the small hill (Fig. 15B). Outcrops in the vicinity of Charleston Reservoir provide an outstanding view of a “dam-burst flood” megabreccia deposit, as well as overlying ash-flow tuff, and sedimentary deposits. Cook and Brueseke (2010) presented more detailed descriptions of the Charleston Reservoir deposits and used major and trace element geochemistry, in conjunction with published data, to correlate the tuffs to other Eocene tuffs in the region. Vitrophyric ash-flow tuff float (Tcm; Fig. 15B) may be apparent in the flat and slopes leading up

59

to the megabreccia outcrop that forms the ridgeline. Geochemical analyses of the vitrophyre indicates that it is the Tuff of Coal Mine Canyon (Cook and Brueseke; 2010), which had not previously been identified in this region. Its presence here and to the north in Copper Basin provides evidence of a paleovalley intersecting the north paleovalley of Henry (2008) and is consistent with a source caldera located to the southwest. The megabreccia exposure is ~4.5 m high (Fig. 15C). This deposit forms a resistant topographic high of semi-continuous outcrop between the two roads (Fig. 15B). Breccia clasts are angular to subrounded, pebble to boulder-sized, plagioclase-biotite tuff (Cook and Brueseke, 2010; Fig. 15D; 0624776 4604533). The matrix consists of medium to coarse tuffaceous sand and gravel and sparse, <1 cm fragments of Paleozoic limestone. Although this breccia resembles caldera mesobreccia, its location and contact relationships to adjacent units suggest that its origin was via damburst type flood events, like other paleovally megabreccias studied by Henry (2008). These occur after ash-flow tuff deposition blocks a drainage and the tuff “dam” catastrophically fails (Henry, 2008; see also Stop 3-1). Walk/scramble up through the exposure and continue in a northwest direction. For the next ~1/3 mi, plagioclase-biotite tuff boulders are sporadically exposed along the upper surface of the megabreccia. After walking up and over a small rise, look for tuff of Coal Mine Canyon vitrophyre float in the topographic low (Cook and Brueseke, 2010; Fig. 15B; 0624468 4604778). This vitrophyre is chemically identical (thus stratigraphically equivalent), to the tuff of Coal Mine Canyon vitrophyre that crops out between the well-exposed megabreccia outcrops and the parking area (Cook and Brueseke, 2010). Cook and Brueseke (2010) suggest that this ash-flow tuff flowed down the same paleovalley as the plagioclase-biotite tuff that forms the megabreccia but flowed around and over the megabreccia. This paleovalley is interpreted to have extended to the north to Copper Basin, on the basis of a newly identified outcrop of the tuff of Coal Mine Canyon in Copper Basin (Cook and Brueseke, 2010). The tuff of Coal Mine Canyon that overlies the megabreccia dips ≤10° NW. Walk another 20–30 ft northwest and move out of the basal vitrophyre into non-vitrophyric ash-flow tuff (0624371 4604789). Continue walking northwest to the small ridge directly ahead and encounter a new clastic unit (Tbx; Fig. 15B; 0624228 4604821). This unit contains angular to subrounded clasts of limestone, ash-flow tuff (including vitrophyre), and other lithics and is silicified in places. Based on mapping and aerial image interpretation, it overlies the tuff of Coal Mine Canyon (also exposed to the SE, adjacent to tan silt deposits; 0624012 4604639). The age of the silt deposits is unclear; however this unit overlies the Eocene deposits and is likely part of the younger package of sediments that fill the basin to the north and south (Humboldt Formation?; Coats, 1987). Stop 2-4 (Optional). Eocene Ash-Flow Tuff Stratigraphy Please watch for rattlesnakes. This stop is along Elko County Road 746 ~9.1 mi west of Charleston Reservoir (11.7 mi east of NV 225 [Mountain City

C

D

Tts: tuffaceous sedimentary rocks Tcm: tuff of Coal Mine Canyon; subscript v indicates vitrophyre Tbx: conglomerate/breccia Tpb: plagioclase-biotite tuff fluvial breccia

B

Figure 15. (A) Oxidized breccia on Jarbidge Rhyolite flow-top, north of Charleston Reservoir. Hammer is 55 cm long. (B) Complex relationships between fluvial megabreccia and overlying ash-flow tuff and sediments at Charleston Reservoir. Tcm lies above (and dips to the NW) a topographic high of Tpb. Tcmv exposed SE of the Tpb megabreccia is chemically identical (thus stratigraphically equivalent) to Tcm exposed NW of megabreccia. Dashed lines are inferred contacts and solid line is contact. Oblique Google Earth view simplified from Cook and Brueseke (2010). (C) Tpb fluvial megabreccia at Charleston Reservoir. Person is ~1.7 m tall. View to west. (D) Close-up of Tpb fluvial megabreccia. Notice the wide range in clast size and clast with tan fiamme (arrow) to the right of the hammer (~50 cm long).

A

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Cenozoic extension in the northern Great Basin Highway]) where a dirt track leads to the southeast (612380 4602076). This stop provides an overview of ash-flow tuff stratigraphy (Tt1; Coats, 1987) and subsequent faulting that characterizes the region between Charleston Reservoir and NV 225. The cliffs just south of the road east of this point are plagioclasebiotite tuff, while the parking area sits on the dipslope of the southeast-dipping 45 Ma tuff, which underlies the plagioclase biotite tuff (Cook and Brueseke, 2010). If time allows, walk northwest to observe excellent exposures of 45 Ma tuff (0612183 4602201). North-northeast–trending normal faults in this region cut the ash-flow tuffs, leading to complex stratigraphic relationships between Eocene eruptive units and the paleovalley inferred by Cook and Brueseke (2010) that goes through the Charleston Reservoir region toward Copper Basin. Detailed mapping is needed to further refine these relationships. Day 3. Eastern Continuation of Central Paleovalley Take the Nevada Highway 233 (Oasis) exit off I-80, turn south on the good gravel road and drive ~1.5 mi south to a side road to the northwest. Take this road ~1.9 mi to a junction and take the right-hand fork about another 0.8 mi to whatever convenient parking can be found (Fig. 16). The road continues and can be driven to some of the initial points on the traverse below. Stop 3-1. Nanny Creek Paleovalley, Ash-flow Tuffs, and Fluvial Megabreccia Nanny Creek, one of the best-exposed and most informative paleovalleys in northeastern Nevada, illustrates paleovalley geometry, enclosed sedimentary deposits, distal ash-flow tuffs, and fluvial megabreccia consisting of reworked ashflow tuff (Fig. 16; Henry, 2008; Brooks et al., 1995a, 1995b). The stratigraphic section is exposed along an ~5 km, nonstrenuous traverse. The exposed width of the Nanny Creek paleovalley is ~6 km, but it is faulted against Paleozoic rocks on the south side and the original width is thus unknown. Mississippian and Permian sedimentary rocks make up the paleovalley wall (Thorman, 1970; Brooks et al., 1995a, 1995b; Camilleri, 2010). A 20–30-m-thick, basal Tertiary conglomerate consists of a lag of rounded boulders commonly up to 1.5 m in diameter, and with one 6 m across (706009, 4544578). Most clasts, including the largest ones, are chert ± quartzite-pebble conglomerate similar to rocks of the Chainman, Diamond Peak, or Dale Canyon Formations, which crop out in immediately adjacent paleovalley walls (Thorman, 1970; Camilleri, 2010). So the largest clasts need not have been transported far. However, numerous clasts of coarse-grained granite up to 1.7 m in diameter, andesite, and silicified, finely laminated lake-bed sediments are present, rock types that do not crop out locally (Brooks et al., 1995a). A granite boulder that we will see (706045, 4544464) has a U-Pb zircon rumor-chron date of ca. 158 Ma (Top Secret Data, 2010). Unfortunately, bedding to determine dip in the conglomerate is not exposed.

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Conglomerate is overlain by a plagioclase-biotite tuff, dated at 40.7 Ma (706115, 4544488), which is in turn overlain by the 40.0 Ma tuff of Big Cottonwood Canyon (706329, 4544541; Fig. 16). The plagioclase-biotite tuff is petrographically and compositionally similar to, but possibly younger than, a plagioclase-biotite tuff in Copper Basin. Tuff thicknesses are uncertain because of uncertainty in their dip. Foliation in the tuffs dips 30–60° eastward, but dips are probably partly primary, resulting from compaction of the tuffs against paleovalley walls (Henry, 2008; Henry and Faulds, 2010). Using the 30–60° range, the plagioclase-biotite tuff and tuff of Big Cottonwood Canyon could be 60–140 m and 90–190 m thick, respectively. Both values are significantly greater than the thicknesses of 25 m and 55 m, respectively, reported by Brooks et al. (1995a) from this same locality. The tuff of Big Cottonwood Canyon here is 150 km from its source in the Tuscarora volcanic field, yet it is at least 55 m thick and restricted to a paleovalley within which it wedges out against the sides. Andesite and andesite flow breccia dated at 39.5 ± 1.0 Ma (Brooks et al., 1995a, 1995b) overlie the tuff of Big Cottonwood Canyon. The uppermost unit in Nanny Creek is a breccia (Tx, Fig. 16) composed of blocks, mostly of silicified tuff of Big Cottonwood Canyon with lesser plagioclase-biotite tuff and andesite (from 706799, 4544865 through 708236, 4544944). PreCenozoic clasts are rare (708016, 4544946). Clasts of tuff of Big Cottonwood Canyon are up to 12 m across (707463, 4545044), plagioclase-biotite tuff is up to 8 m long (707459, 4545061), and andesite is as much as 2 m across (708137, 4544789). Blocks range from angular to rounded, and many are internally broken but not disaggregated (707975, 4544954). Rarely exposed matrix consists of coarse sand and variably rounded pebbles and cobbles. Brooks et al. (1995a) recognized the hummocky character of this unit and the possibility that the hummocks might be landslide blocks. Chemical analysis and dating of one clast confirm that it is tuff of Big Cottonwood Canyon (Henry, 2008). Furthermore, highly discordant compaction foliations and scattered magnetization directions from several breccia clasts (M.R. Hudson, in Brooks et al., 1995a; Palmer and MacDonald, 2002, their sites W02, W15, and W18) confirm that these blocks are not in place. We interpret these breccia deposits as resulting from “dam-burst floods,” in which the tuff of Big Cottonwood Canyon initially blocked a drainage, then the tuff “dam” collapsed once a lake backed up behind and overtopped the dam. Blockage of drainages and dam-burst floods are common even in historic times (Costa and Schuster, 1988; Waythomas, 2001; the Thistle landslide that blocked the Spanish Fork River in Utah in 1982). ACKNOWLEDGMENTS We greatly appreciate discussions with David John, Keith Howard, Chuck Thorman, Alan Wallace, Chuck Chapin, Dave Boden, Jim Faulds, Patrick Goldstrand, John Muntean, Fred Zoerner, Ken Hickey, Dick Tosdal, and Mike Ressel, and very helpful reviews by Keith Howard and Jeffrey Lee.

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40Ar/39Ar date (regular type, Henry, 2008;

114°32"

Tbcc

italics, Brooks et al., 1995a, b)

Qa Eocene Tx

41°3'

Tbcc

Tbcc

Alluvium, alluvial fan deposits

Qf

Ta

Qa Tbcc

Pz Megabreccia

Ta Tbcc

Ta Tbcc

Andesite Tuff of Big Cottonwood Canyon

Tpb

Plagioclase-biotite tuff

Tog

Conglomerate

Qf

Ta

Qf

Tx

Qa

Tx Ta

Pz Pz

50

47

Paleozoic Rocks Qf Tbcc

41°2' 41°2"

60

35

40.10±0.07

37

Qf Qf

54

35

Tog

Qf

Tx

52

Qa

Tbcc Tpb 40.69±0.13

Pz

Ta

47

Qf

Qf

39.93±0.08

Tx

41.35±0.22

Qa

52 Tog

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39.48±1.00

59

39.87±0.26

Approximate Tx parking and field trip traverse Stop 3-1

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Ta Tbcc Tbcc

Tx

Ta

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60 65

41°1'

41°1"

Tbcc

Qa

Qf Tbcc

N 114°33'

Tpb

Pz

0 0 114°32"

0.5 mi 0.5 km

Figure 16. Geologic map of the paleovalley at Nanny Creek, eastern Pequop Mountains, showing traverse for Stop 3-1. The paleovalley, which was at least 6 km wide and possibly 1.6 km deep, has basal conglomerate containing rounded boulders up to 6 m in diameter, overlain by a plagioclase-biotite tuff and tuff of Big Cottonwood Canyon, andesite lavas, and a thick fluvial megabreccia consisting of angular blocks mostly of the tuff of Big Cottonwood Canyon.

Cenozoic extension in the northern Great Basin REFERENCES CITED Armstrong, R.L., 1968, Sevier orogenic belt in Nevada and Utah: Geological Society of America Bulletin, v. 79, p. 429–458, doi:10.1130/0016 -7606(1968)79[429:SOBINA]2.0.CO;2. Armstrong, R.L., and Hansen, E., 1966, Cordilleran infrastructure in the eastern Great Basin: American Journal of Science, v. 264, p. 112–127, doi:10.2475/ajs.264.2.112. Axelrod, D.I., 1966a, Potassium-Argon ages of some western Tertiary floras: American Journal of Science, v. 264, p. 497–506, doi:10.2475/ ajs.264.7.497. Axelrod, D.I., 1966b, The Eocene Copper Basin flora of northeastern Nevada: University of California Publications in Geological Sciences, v. 59, 124 p. Baars, D.L., Bartleson, B.L., Chapin, C.E., Curtis, B.F., De Voto, R.H., Everett, J.R., Johnson, R.C., Molenaar, C.M., Peterson, F., Schenk, C.J., Love, J.D., Merin, I.S., Rose, P.R., Ryder, R.T., Waechter, N.B., and Woodward, L.A., 1988, Basins of the Rocky Mountain region, in Sloss, L.L., ed., Sedimentary Cover—North American Craton: U.S.: Geological Society of America, Geology of North America, v. D-2, p. 109–220. Barton, M.D., 1996, Granitic magmatism and metallogeny of southwestern North America, in Brown, M., Candela, P.A., Peck, D.L., Stephens, W.E., Walker, R.J., and Zen, E-An, eds., Origin of Granites and Related Rocks: Geological Society of America Special Paper 315, p. 261–280. Best, M.G., and Christiansen, E.H., 1991, Limited extension during peak Tertiary volcanism, Great Basin of Nevada and Utah: Journal of Geophysical Research, v. 96, p. 13,509–13,528, doi:10.1029/91JB00244. Best, M.G., Barr, D.L., Christiansen, E.H., Gromme, S., Deino, A.L., and Tingey, D.G., 2009, The Great Basin Altiplano during the middle Cenozoic ignimbrite flareup: insights from volcanic rocks: International Geology Review, v. 51, p. 589–633, doi:10.1080/00206810902867690. Bohlen, S.R., Montana, A., and Kerrick, D.M., 1991, Precise determination of the equilibria kyanite-sillimanite and kyanite-andalusite and revised triple point for Al2SiO5 polymorphs: The American Mineralogist, v. 76, p. 677–680. Brooks, W.E., Thorman, W.E., and Snee, L.W., 1995a, The 40Ar/39Ar ages and tectonic setting of the middle Eocene northeast Nevada volcanic field: Journal of Geophysical Research, v. 100, p. 10,403–10,416, doi:10.1029/94JB03389. Brooks, W.E., Thorman, W.E., Snee, L.W., Nutt, C.W., Potter, C.J., and Dubiel, R.F., 1995b, Summary of chemical analyses and 40Ar/39Ar-spectra data for Eocene volcanic rocks from the central part of the northeast Nevada volcanic field: U.S. Geological Survey Bulletin 1988-K, p. K1–K33. Bryant, B., Naeser, C.W., Marvin, R.F., and Mehnert, H.H., 1989, Upper Cretaceous and Paleogene sedimentary rocks and isotopic ages of Paleogene tuffs, Uinta Basin, Utah: U.S. Geological Survey Bulletin 1787-J, 22 p. Bushnell, K., 1967, Geology of the Rowland Quadrangle, Elko County, Nevada: Nevada Bureau of Mines Bulletin 67, 38 p. Callicoat, J.S., 2010, Significance of mid-Miocene volcanism in northeast Nevada: petrographic, chemical, isotopic, and temporal importance of the Jarbidge Rhyolite [M.S. thesis]: Manhattan, Kansas, Kansas State University, p. 1–108. Callicoat, J., Brueseke, M., and Larson, P.B., 2010, Oxygen isotope constraints on voluminous mid-Miocene effusive silicic magmatism in north-central Nevada: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 101. Camilleri, P., 2010, Geologic map of the Wood Hills, Elko County, Nevada: Nevada Bureau of Mines and Geology Map 172, 1:48,000. Camilleri, P., and Chamberlain, K.R., 1997, Mesozoic tectonics and metamorphism in the Pequop Mountains and Wood Hills, northeast Nevada: Implications for the architecture and evolution of the Sevier orogen: Geological Society of America Bulletin, v. 109, p. 74–94, doi:10.1130/0016 -7606(1997)109<0074:MTAMIT>2.3.CO;2. Camilleri, P., Yonkee, W.A., Coogan, J.C., DeCelles, P.G., McGrew, A., and Wells, M., 1997, Hinterland to foreland transect through the Sevier orogen, NE Nevada to SW Wyoming: structural style, metamorphism, and kinematic history of a large contractional orogenic wedge, in Link, P.K., and Kowallis, B.J., eds., Proterozoic to Recent Stratigraphy, Tectonics, and Volcanology, Utah, Nevada, Southern Idaho and Central Mexico: Brigham Young University Geology Studies, v. 42, part 1, p. 297–309. Cassel, E.J., Henry, C.D., Graham, S.A., and Chamberlain, P.C., 2010, Determining Oligocene topography and tectonism across the northern Sierra Nevada and western Basin and Range using stable isotope paleoaltimetry in volcanic glass: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 184.

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Chase, C.G., Gregory, K.M., Parrish, J.T., and DeCelles, P.G., 1998, Topographic history of the western Cordilleran of North America and the etiology of climate, in Crowley, T.J., and Burke, K., eds., Tectonic boundary conditions for climate reconstructions: Oxford Monographs on Geology and Geophysics, no. 39, p. 73–99. Christiansen, R.L., and Yeats, R.S., 1992, Post-Laramide geology of the U.S. Cordilleran region, in Burchfiel, B.C., Lipman, P.W., and Zoback, M.L., eds., The Cordilleran Orogen: Conterminous U.S.: Geological Society of America, Geology of North America, v. G-3, p. 261–406. Cline, J.S., Hofstra, A.H., Muntean, J.L., Tosdal, R.M., and Hickey, K.A., 2005, Carlin-type gold deposits in Nevada: Critical geologic characteristics and viable models, in Hedenquist, J.W., Thompson, J.F.H., Goldfard, R.J., and Richards, J.P., eds., 100th Anniversary Volume, Economic Geology, p. 451–484. Coash, J.R., 1967, Geology of the Mount Velma Quadrangle, Elko County, Nevada: Nevada Bureau of Mines Bulletin 68, 20 p. Coats, R.R., 1964, Geology of the Jarbidge Quadrangle, Nevada–Idaho: U.S. Geological Survey Bulletin 1141-M, 24 p. Coats, R.R., 1987, Geology of Elko County, Nevada: Nevada Bureau of Mines and Geology Bulletin 101, 112 p. Coats, R.R., and Riva, J.F., 1983, Overlapping thrust belts of late Paleozoic and Mesozoic ages, northern Elko County, Nevada, in, Miller, D.M., Todd, V.R., and Howard, K.A., eds., Tectonic and Stratigraphic Studies in the eastern Great Basin: Geological Society of America Memoir 157, p. 305–327. Colgan, J.P., and Henry, C.D., 2009, Rapid middle Miocene collapse of the Sevier orogenic plateau in north-central Nevada: International Geology Review, v. 51, p. 920–961, doi:10.1080/00206810903056731. Colgan, J.P., John, D.A., Henry, C.D., and Fleck, R.J., 2008, Large-magnitude Miocene extension of the Caetano caldera, Shoshone and Toiyabe Ranges, Nevada: Geosphere, v. 4, p. 107–131, doi:10.1130/GES00115.1. Colgan, J.P., Howard, K.A., Fleck, R.J., and Wooden, J.L., 2010, Rapid middle Miocene extension and unroofing of the southern Ruby Mountains, Nevada: Tectonics, v. 29, p. TC6022, doi:10.1029/2009TC002655. Coney, P.J., and Harms, T.A., 1984, Cordilleran metamorphic core complexes: Cenozoic extensional relics of Mesozoic compression: Geology, v. 12, p. 550–554, doi:10.1130/0091-7613(1984)12<550:CMCCCE>2.0.CO;2. Cook, C.C., and Brueseke, M.E., 2010, Petrography and identification of Eocene ash-flow tuffs in the vicinity of the Jarbidge Mountains, Nevada: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 285. Costa, J.E., and Schuster, R.L., 1988, The formation and failure of natural dams: Geological Society of America Bulletin, v. 100, p. 1054–1068, doi:10.1130/0016-7606(1988)100<1054:TFAFON>2.3.CO;2. Crafford, A.E.J., 2007, Geologic map of Nevada: U.S. Geological Survey Data Series 249, 46 p. Dallmeyer, R.D., Snoke, A.W., and McKee, E.H., 1986, The MesozoicCenozoic tectonothermal evolution of the Ruby Mountains, East Humboldt Range, Nevada: a Cordilleran metamorphic core complex: Tectonics, v. 5, p. 931–954, doi:10.1029/TC005i006p00931. Davis, S.J., Mulch, A., Carroll, A.R., Horton, T.W., and Chamberlain, C.P., 2009, Paleogene landscape evolution of the central North American Cordillera: Developing topography and hydrology in the Laramide foreland: Geological Society of America Bulletin, v. 121, p. 100–116, doi:10.1130/B26308.1. DeCelles, P.G., 2004, Late Jurassic to Eocene evolution of the Cordilleran thrust belt and foreland basin system, western U.S.A: American Journal of Science, v. 304, p. 105–168, doi:10.2475/ajs.304.2.105. DeCelles, P.G., and Coogan, J.C., 2006, Regional structure and kinematic history of the Sevier fold and-thrust belt, central Utah: Geological Society of America Bulletin, v. 118, p. 841–864, doi:10.1130/B25759.1. dePolo, C.M., Smith, K.D., and Henry, C.D., 2011, Summary of the 2008 Wells, Nevada earthquake documentation volume, in dePolo, C.M., and LaPointe, D.D., eds., A compendium of earthquake-related investigations prepared by the University of Nevada, Reno: Nevada Bureau of Mines and Geology Special Publication 36, p. 7–14. Dickinson, W.R., 2002, The Basin and Range Province as a composite extensional domain: International Geology Review, v. 44, p. 1–38, doi:10.2747/0020-6814.44.1.1. Dickinson, W.R., 2006, Geotectonic evolution of the Great Basin: Geosphere, v. 2, p. 353–368, doi:10.1130/GES00054.1. Dilek, Y., and Moores, E.M., 1999, A Tibetan model for the early Tertiary western United States: Journal of the Geological Society, v. 156, p. 929–941, doi:10.1144/gsjgs.156.5.0929. Dokka, R.K., Mahaffie, M.J., and Snoke, A.W., 1986, Thermochronologic evidence of major tectonic denudation associated with detachment faulting,

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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. Thorman, C.H., 1970, Metamorphosed and nonmetamorphosed Paleozoic rocks in the Wood Hills and Pequop Mountains, northeast Nevada: Geological Society of America Bulletin, v. 81, p. 2417–2448, doi:10.1130/0016 -7606(1970)81[2417:MANPRI]2.0.CO;2. Thorman, C.H., Ketner, K.B., Brooks, W.E., Snee, L.W., and Zimmerman, R.A., 1991, Late Mesozoic–Cenozoic tectonics in northeastern Nevada, in Raines, G.L., Lisle, R.E., Schafer, R.W., and Wilkinson, W.H., eds., Geology and Ore Deposits of the Great Basin Symposium: Geological Society of Nevada, Reno, Proceedings, v. 1, p. 25–45. Vanderhaeghe, O., Teyssier, C., and Ord, A., eds., 2001, Crustal-scale rheological transitions during late-orogenic collapse: Tectonophysics, v. 335, p. 211–228, doi:10.1016/S0040-1951(01)00053-1. Vandervoort, D.S., and Schmitt, J.G., 1990, Cretaceous to early Tertiary paleogeography in the hinterland of the Sevier thrust belt, east-central Nevada: Geology, v. 18, p. 567–570, doi:10.1130/0091-7613(1990)018<0567: CTETPI>2.3.CO;2. Wallace, A.R., Perkins, M.E., and Fleck, R.J., 2008, Late Cenozoic paleogeographic evolution of northeastern Nevada: Evidence from the sedimentary basins: Geosphere, v. 4, p. 36–74, doi:10.1130/GES00114.1. Waythomas, C.F., 2001, Formation and failure of volcanic debris dams in the Chakachatna River valley associated with eruptions of the Spurr volcanic complex, Alaska: Geomorphology, v. 39, p. 111–129, doi:10.1016/S0169 -555X(00)00097-0. Wells, M.L., and Hoisch, T.D., 2008, The role of mantle delamination in widespread Late Cretaceous extension and magmatism in the Cordilleran Orogen, western United States: Geological Society of America Bulletin, v. 120, p. 515–530, doi:10.1130/B26006.1. Wernicke, B., 1992, Cenozoic extensional tectonics of the U.S. Cordillera, in Burchfiel, B.C., Lipman, P.W., and Zoback, M.L., eds., The Cordilleran Orogen: Conterminous U.S.: Geological Society of America, Geology of North America, v. G-3, p. 553–581. Wernicke, B.P., and Getty, S.R., 1997, Intracrustal subduction and gravity currents in the deep crust: Sm-Nd, Ar-Ar, and thermobarometric constraints from the Skagit Gneiss Complex, Washington: Geological Society of America Bulletin, v. 109, p. 1149–1166, doi:10.1130/0016 -7606(1997)109<1149:ISAGCI>2.3.CO;2. Wickham, S.M., and Peters, M.T., 1992, Oxygen and carbon isotope profiles in metasediments from Lizzies Basin, East Humboldt Range, Nevada; constraints on mid-crustal metamorphic and magmatic volatile fluxes: Contributions to Mineralogy and Petrology, v. 112, p. 46–65, doi:10.1007/BF00310955. Wolfe, J.A., Forest, C.E., and Molnar, P., 1998, Paleobotanical evidence of Eocene and Oligocene paleoaltitudes in midlatitude western North America: Geological Society of America Bulletin, v. 110, p. 664–678, doi:10.1130/0016-7606(1998)110<0664:PEOEAO>2.3.CO;2. Wooden, J.L., Kistler, R.W., Tosdal, R.M., Robinson, A., and Wright, J.E., 1997, Pb vs. Sr isotopic mapping of crustal structure in the northern Great Basin: Geological Society of America Abstracts with Programs, v. 29, p. 70. Woods, A.W., Bursik, M.I., and Kurbatov, A.V., 1998, The interaction of ash flows with ridges: Bulletin of Volcanology, v. 60, p. 38–51, doi:10.1007/ s004450050215. Wright, J.E., and Snoke, A.W., 1993, Tertiary magmatism and mylonitization in the Ruby–East Humboldt metamorphic core complex, northeastern Nevada: U-Pb geochronology and Sr, Nd, and Pb isotope geochemistry: Geological Society of America Bulletin, v. 105, p. 935–952, doi:10.1130/0016-7606(1993)105<0935:TMAMIT>2.3.CO;2. Wright, J.E., and Wooden, J.L., 1991, New Sr, Nd, and Pb isotopic data from plutons in the northern Great Basin: Implications for crustal structure and granite petrogenesis in the hinterland of the Sevier thrust belt: Geology, v. 19, p. 457–460, doi:10.1130/0091-7613(1991)019<0457:NSNAPI>2.3.CO;2. Zoback, M.L., McKee, E.H., Blakely, R.J., and Thompson, G.A., 1994, The northern Nevada rift: Regional tectonomagmatic relations and middle Miocene stress direction: Geological Society of America Bulletin, v. 106, p. 371–382, doi:10.1130/0016-7606(1994)106<0371:TNNRRT>2.3.CO;2. MANUSCRIPT ACCEPTED BY THE SOCIETY 4 MARCH 2011 Printed in the USA

The Geological Society of America Field Guide 21 2011

Tectonomagmatic evolution of distinct arc terranes in the Blue Mountains Province, Oregon and Idaho C.J. Northrup M. Schmitz G. Kurz K. Tumpane Department of Geosciences, Boise State University, 1910 University Drive, Boise, Idaho 83725, USA

ABSTRACT Recent mapping, U-Pb zircon geochronology, trace-element geochemistry, and tracer isotope geochemistry of plutonic and volcanic rocks in the Wallowa and Olds Ferry terranes of the Blue Mountains Province yield new insights into their tectonic evolution and pre-accretion history. Igneous rocks of the Wallowa arc terrane formed in two magmatic episodes of contrasting duration and geochemical characteristics. Magmatism in the first episode lasted for at least 20 Ma (ca. 268–248 Ma), spanning the Middle Permian to the Early Triassic and was of generally calc-alkaline affinity. Rock units associated with this episode include the Hunsaker Creek and Windy Ridge formations of the Wallowa terrane, as well as potentially equivalent tonalite and diorite plutonic rocks in the Cougar Creek Complex and related basement exposures, which show midcrustal levels of the terrane. The second episode of magmatism in the Wallowa arc was remarkably brief (U-Pb zircon dates range from 229.43 ± 0.08 Ma to 229.13 ± 0.45 Ma) and dominated by mafic to intermediate compositions of tholeiitic affinity. Rock units associated with the second episode may include the Wild Sheep Creek and Doyle Creek formations, as well as ubiquitous dikes and plutons in the Cougar Creek Complex and similar basement exposures. After 229 Ma, the Wallowa arc apparently became dormant. The record of igneous activity in the Olds Ferry arc contrasts with that of the Wallowa in its age range and the continuity of calc-alkaline magmatism. Radiometric ages and stratigraphic field relationships allow the magmatic history of the Olds Ferry terrane to be divided into at least three cycles separated by brief hiatuses and collectively spanning the late Middle Triassic through the Early Jurassic (ca. 237– 187 Ma). Rock units related to these episodes are divided by unconformities, and they include the Brownlee pluton, lower Huntington Formation, and upper Huntington Formation. Magmatic activity in the Olds Ferry arc may have persisted until at least 174 Ma, based on the presence of volcanic ash horizons in the lower portion of the overlying Weatherby Formation of the Izee basin. All cycles of Olds Ferry magmatism display generally calc-alkaline affinity.

Northrup, C.J., Schmitz, M., Kurz, G., and Tumpane, K., 2011, Tectonomagmatic evolution of distinct arc terranes in the Blue Mountains Province, Oregon and Idaho, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 67–88, doi: 10.1130/2011.0021(03). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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Northrup et al. The contrasting magmatic histories of the Wallowa and Olds Ferry arc terranes provide the basis for at least two conclusions. First, these arcs formed as separate tectonic entities, rather than as a single composite arc. Second, progressive closure of the ocean basin between the arcs in the Late Triassic and Early Jurassic was related to continued subduction beneath the Olds Ferry arc, but the Wallowa arc was apparently dormant during much of that interval.

INTRODUCTION The timing, paleogeography, and kinematic history of late Paleozoic–early Mesozoic terrane accretion along the western margin of North America remain fundamental issues in the assembly and evolution of the Cordilleran orogen. Throughout much of the northwestern United States, the products of terrane accretion and crustal assembly are obscured by an extensive blanket of younger cover, including voluminous volcanic rocks of the Columbia River Basalts (Fitzgerald, 1982; Swanson et al., 1981). However, in the Blue Mountains province of IdahoOregon-Washington, regional uplift and incision of the Snake River and its tributaries through the Cenozoic cover have gener-

ated spectacular exposures of the accreted Paleozoic and Mesozoic “basement” (Figs. 1 and 2; Brooks et al., 1976). Pre-Cenozoic rocks of the Blue Mountains province record a protracted and complex tectonic history, spanning the Permian to the Cretaceous (Brooks and Vallier, 1978; Dickinson, 1979; Avé Lallemant et al., 1980; Wilson and Cox, 1980, Hillhouse et al., 1982; White et al., 1992; Vallier, 1995; Gray and Oldow, 2005; Dorsey and LaMaskin, 2007). Previous studies have identified at least four first-order tectonic elements (terranes) in the Blue Mountains province, outboard of the ancient cratonal margin of North America as defined by initial 87Sr/86Sr isotopic ratios of Mesozoic granitic plutons in the region (Hamilton, 1963; Davis et al., 1978; Dickinson, 1979; Dickinson and Thayer, 1978;

Figure 1. Tectonic map of the western North American Cordillera showing arc assemblages in present day (A) and before ~60° clockwise rotation of Blue Mountains province and Oregon-Washington Coast Range (B). Arc assemblages: In— Insular (SG—Swakane Gneiss; W-SD—Wallowa–Seven Devils); Km—accreted western Klamath Mountains arcs; Qu— Quesnellia and related terranes (IZ—Izee forearc basin; EK—eastern Klamath Mountains terrane; OF—Olds Ferry terrane); St—Stikinia (CR-H—Cascade River–Holden belt). Accretionary prism/subduction complex terranes: B—Baker; BR—Bridge River; CC—Cache Creek; H—Hozameen; K—central Klamath Mountains mélange belt. Other features: Sh—Shuksan thrust system (schematic); TMt—Tyaughton-Methow trough (offset segments: Mt—Methow trough; Tt— Tyaughton trough); SCf—Straight Creek fault; RLf—Ross Lake fault; FRf—Fraser River fault; Yf—Yalakom fault. From Dickinson (2004).

Blue Mountains Province, Oregon and Idaho Armstrong et al., 1977). The Wallowa and Olds Ferry terranes contain Permian and Triassic plutonic, volcanic, and volcaniclastic lithologies overlain by Upper Triassic to Jurassic volcanic, volcaniclastic, and sedimentary rocks. The Baker terrane is a mélange complex of deformed and metamorphosed Triassic and Jurassic argillite and chert, with olistrostromal blocks of Devonian to Permian limestone, formed by marine sedimentation followed by intense deformation in an oceanic subduction zone (Schwartz et al., 2010). The Izee terrane consists of Upper Triassic to Upper Jurassic sedimentary rocks with a wide range of ages, lithofacies, and depositional environments (dominantly shallow to deep marine) (Dickinson, 1979; Dorsey and LaMaskin, 2007). These strata depositionally overlie rocks of the Wallowa, Olds Ferry, and Baker terranes, and thus may represent an overlap assemblage that links the three terranes by the late Middle Jurassic (Pessagno and Blome, 1986; White et al., 1992; Dorsey and LaMaskin, 2007). Late Jurassic to Early Cretaceous plutons cut all earlier units, structures, and fabrics, but are themselves variably affected by the continued tectonic evolution of the region. Arc-related terranes of the Blue Mountains province occupy a critical position within the Cordilleran Orogen, providing a link between better documented areas of Mesozoic terrane accretion to the south in the Klamath Mountains, and north in British Columbia and Alaska (Fig. 1; Jones et al., 1977; Davis et al. 1978; Saleeby, 1983; Silberling et al., 1984; Burchfiel et al., 1992; Dickinson, 2004). A comprehensive view of crustal assembly and terrane correlation along the western margin of North

Figure 2. Geologic map of the Blue Mountains province showing the four major terranes as well as the extensive Cenozoic cover. General areas examined in this field trip are near Huntington (H, Stops 1.1– 1.6), and Pittsburg Landing (P, Stops 2.1–2.3). IB—Idaho Batholith, BC—Baker City, WISZ—Western Idaho Shear Zone. Modified from Vallier (1995).

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America requires a detailed understanding of the evolution of the Blue Mountains province terranes and how they compare with terranes in other locations. The relationship between the Olds Ferry and Wallowa terranes has been an important, recurrent theme in studies of the Blue Mountains province and its tectonic history. Some authors have suggested that both terranes were components of a single, larger Blue Mountains island arc (Charvet et al., 1990; Pessagno and Blome, 1986; Vallier and Brooks, 1986; Vallier and Engebretson, 1983; White et al., 1992), while others have suggested that the volcanic arcs evolved separately for a period of time prior to amalgamation and accretion to North America (Saleeby et al., 1992; Avé Lallemant, 1995; Vallier, 1995; Dorsey and LaMaskin, 2007; LaMaskin et al., 2008a; Schwartz et al., 2010). The purpose of this excursion is to examine key field relationships and discuss the results of recent high-precision U-Pb geochronology and tracer isotope geochemistry that help to clarify and distinguish the tectonic histories of the Old Ferry and Wallowa arc terranes. We compare and contrast the volcano-plutonic records of these arc systems, evaluate models of their tectonic and paleogeographic evolution from the upper Paleozoic through the Early Cretaceous, and highlight possible correlations of the Blue Mountains arc terranes with other tectonic elements exposed along the Cordilleran margin. Field trip stops and discussion will focus primarily on the geology of the Olds Ferry terrane on Day 1 and the Wallowa terrane on Day 2. GEOLOGY OF THE OLDS FERRY TERRANE The Olds Ferry terrane contains the most inboard expression of Permo-Triassic arc volcanism in the Blue Mountains with respect to the North American craton. It lies east of the subduction mélange complex of the Baker terrane, which is one of the “Cache Creek affinity” terranes distributed along the cordilleran margin of North America (Burchfiel et al., 1992; Kays et al., 2006). This position suggests correlation with other marginal magmatic arc rocks along the Cordilleran orogen like the Quesnel and Eastern Klamath Mountains terranes (Dickinson, 2004; Dorsey and LaMaskin, 2007; LaMaskin et al., 2008a; Schwartz et al., 2010). The most extensive exposures of plutonic, volcanic, volcaniclastic, and sedimentary rocks of the Olds Ferry terrane are found along the Snake River canyon, beginning near the confluence of the Burnt and Snake rivers east of Huntington, Oregon, and continuing northward for ~11 km to the Bay Horse mine (Fig. 3). These exposures include the Brownlee pluton, the lower Huntington Formation, the upper Huntington Formation, and the contact with the overlying Weatherby Formation of the Izee terrane. Recent U-Pb zircon geochronology of igneous rocks in the Olds Ferry terrane provides a useful framework of age information and helps to delineate the timing of magmatism recorded in this arc terrane (Table 1; Tumpane, 2010; Kurz, 2010). Although the Brownlee pluton is exposed only in a relatively small area (~1.5 km2) spanning both sides of the Snake River, it

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TABLE 1. RECENT U-Pb ZIRCON GEOCHRONOLOGY OF IGNEOUS ROCKS, OLDS FERRY ARC 206 238 Sample Geologic Rock Pb/ U Age identification unit type (Ma) DC 07-01 DC 07-03

Lower Weatherby Formation Lower Weatherby Formation

Welded tuff Crystal tuff

173.91 ± 0.07 180.61 ± 0.17

DC 07-05 DC 08-04

Upper Huntington Formation Upper Huntington Formation

Rhyolite tuff Rhyodacite

187.03 ± 0.04 188.45 ± 0.05

DC 07-06 CUD09-01 RC07-04 HT 05-11 HT 09-03 RC07-03 RC07-06 HT 08-14

Iron Mountain pluton Boulder in Jurassic conglomerate Rush Creek granodiorite Lower Huntington Formation Lower Huntington Formation Rush Creek diorite Rush Creek gabbro Lower Huntington Formation

Hbl granodiorite Quartz diorite Porphyritic granodiorite Dacite tuff breccia Rhyodacite flow Quartz diorite Gabbro Lithic tuff

210.04 ± 0.12 216.48 ± 0.06 219.10 ± 0.10 220.66 ± 0.18 220.52 ± 0.08 220.29 ± 0.06 220.88 ± 0.05 221.72 ± 0.12

HT07-02 Small trondhjemite pluton Trondhjemite 237.60 ± 0.05 HT04-04 Brownlee pluton Trondhjemite 237.68 ± 0.07 Note: All dates are from U-Pb zircon thermal ionization mass spectrometry analyses of Kurz (2010), and Tumpane (2010); See these sources for detailed sample locations, complete data tables, and analytical methods.

is significant because it is the oldest known component of the Olds Ferry terrane, dated at 237.68 ± 0.07 Ma (Kurz, 2010). The pluton consists of trondhjemite, and it has been cut by numerous mafic dikes that likely represent part of the feeder system for volcanic rocks in the overlying Huntington Formation. Contact relationships between the Brownlee pluton and adjacent supracrustal rocks play an important role in illuminating the tectonostratigraphic architecture of the Olds Ferry terrane (Tumpane, 2010), and will be examined in detail at Stop 1.3 of the field trip. The Huntington Formation was formally named by Brooks (1979a) and, along with plutonic rocks of the Cuddy Mountains– Sturgill Peak area to the northeast, comprises the majority of the Olds Ferry terrane. It is a thick succession consisting primarily of volcanic deposits ranging in composition from basalt to rhyolite, but dominated by andesite and basaltic andesite (Brooks, 1979a, 1979b; Brooks et al., 1976; Brooks and Vallier, 1967, 1978; Charvet et al., 1990; Collins, 2000; Dorsey and LaMaskin, 2007; Tumpane, 2010; Vallier, 1995; Wagner et al., 1963). Sedimentary rocks occur as interbeds and include volcanic sandstones, conglomerates and minor, discontinuous limestones. All rocks in the Huntington Formation have been hydrothermally altered and metamorphosed in the zeolite to lower greenschist facies. The thickness of the Huntington Formation is difficult to determine with precision due to stratigraphic and structural complexity, but estimates range from at least 3000 m to more than 6000 m (Brooks 1979a). The Huntington Formation has been divided informally by some authors into an upper and lower member, based on the

Figure 3. Geologic map of Huntington area, as compiled by Tumpane (2010), including data from Brooks (1979a), Collins (2000), and Juras (1973). Geochronologic data for samples of Tumpane (2010) are summarized in Table 1. CRB—Columbia River Basalt Group.

character of the volcanic package and relative abundance of sedimentary components (Juras, 1973; Collins, 2000; Dorsey and LaMaskin, 2007; LaMaskin, 2008). Recent geochronology and detailed mapping indicate that both the upper and lower members are bounded by unconformities (Tumpane, 2010). The lower member of the Huntington Formation is composed dominantly of massive, mafic to intermediate lava flows, sills, volcanic breccias, and tuff breccias, interstratified with minor components of finegrained volcaniclastic sandstones, limestones, and other clastic sedimentary rocks. Ammonites of the Tropites dilleri and Tropites welleri zones have been found within the lower Huntington Formation, indicating a late Carnian to Norian age (ca. 232 Ma; Smith, 1927; Brooks and Vallier, 1978; LaMaskin, 2008; Ogg et al., 2008). U-Pb zircon geochronology of volcanic rocks within the lower Huntington Formation has yielded ages of 221.72 ± 0.12 Ma and 220.52 ± 0.08 Ma (Table 1; Tumpane, 2010), indicating the depositional interval for this unit continued until at least this time. The upper Huntington Formation is distinguished from the lower member by a greater abundance of sedimentary rocks such as volcanic sandstone, turbidites, and laminated shale. In addition, volcanic rocks of the upper Huntington are generally more silicic compared to those of the lower Huntington (Collins, 2000; Dorsey and LaMaskin, 2007; LaMaskin, 2008). New geochronology of the upper member of the Huntington Formation helps to define its depositional interval and extends Olds Ferry volcanism into the Early Jurassic. The age of the base of the upper Huntington Formation is not well constrained. Part of the upper Huntington section rests unconformably on plutonic rocks dated at 210.04 ± 0.12 Ma near the abandoned mining town of Mineral, Idaho, placing a maximum age on the timing of its deposition (Table 1; Kurz, 2010). The age of the uppermost part of the unit is well constrained by rhyodacite and rhyolite tuffs with U-Pb zircon dates of 188.45 ± 0.05 and 187.03 ± 0.04 Ma, respectively

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(Tumpane, 2010). These ages lie within the Pliensbachian stage of the Early Jurassic. Volcanic rocks in both the upper and lower Huntington Formation are generally of calc-alkaline affinity, consistent with formation in a volcanic arc along a convergent plate boundary (Vallier, 1995; Collins, 2000; Tumpane, 2010). Huntington volcanic rocks have 87Sr/86Sr values ranging from 0.7034 to 0.7057 and εNd values of +6.9 to +3.1, and display a trend of generally less radiogenic εNd and more variable and radiogenic 87Sr/86Sr compositions in the upper member compared to the lower member (Tumpane, 2010). Similar geochemical and isotopic characteristics have been found in the plutonic rocks of the Olds Ferry terrane (Kurz, 2010). The Olds Ferry terrane is bounded to the east by the Salmon River Belt (Gray and Oldow, 2005), and the Western Idaho Shear Zone, a subvertical, lithospheric-scale feature of mid-Cretaceous age that juxtaposes the composite terranes of the Blue Mountains province against the margin of cratonal North America (Giorgis et al., 2007; Giorgis et al., 2008). In addition to the mylonitic rocks of the Western Idaho Shear Zone, the abrupt transition between Precambrian continental lithosphere and relatively juvenile oceanic material accreted to the margin is reflected in an abrupt changes in the initial 87Sr/86Sr isotopic composition of Mesozoic plutons and the lithology of pendants and inclusions within them (Fleck and Criss, 1985, 2007; Giorgis et al., 2005; Lund and Snee, 1988; Manduca et al., 1993; McClelland et al., 2000; Selverstone et al., 1992). The Western Idaho Shear Zone likely accommodated lateral translation of the composite Blue Mountains province terranes along the Cordilleran margin after accretion and prior to ca. 90 Ma, by which time it had become inactive (Giorgis et al., 2008; McClelland et al., 2000). The amount and direction of lateral movement, however, remain poorly constrained. The northwestern margin of the Olds Ferry terrane is delineated by the contact of the Huntington Formation and the Weatherby Formation of the Izee terrane (Fig. 3). This boundary is well exposed along the Snake River corridor near Mineral, Idaho, and along the Brownlee Reservoir north of Huntington, Oregon, at the Bay Horse mine (Stop 1.4). The Olds Ferry–Izee contact has been interpreted as a thrust fault (Avé Lallemant, 1983; Livingston, 1932), an unconformity (Brooks, 1967, 1979a, 1979b; Brooks and Vallier, 1967, 1978), and a shear zone detached along an originally depositional contact (Payne and Northrup, 2003; Dorsey and LaMaskin, 2007). Rock units in the lower Weatherby Formation consist of the McChord Butte conglomerate, Dennett Creek limestone, and Big Hill Shale as defined by Henricksen (1975) and Payne and Northrup (2003). The McChord Butte conglomerate passes stratigraphically upward to the Dennett Creek limestone through a transition interval of ~20 m, characterized by a progressive decrease in the grain size of the clastic material to fine sand and an increase in the abundance of interstratified carbonate, culminating in the massive Dennett Creek limestone. The combined thickness of the McChord Butte conglomerate and Dennett

Creek limestone varies significantly (particularly in the thickness of the clastic rocks), but is typically no more than 75 m. The upper contact of the Dennett Creek limestone is sharp and overlain by the Big Hill shale—a thick accumulation of siltstone and mudstone with sporadic tuff and volcaniclastic sandstone. The Big Hill shale is the volumetrically most significant component of the Weatherby Formation. Its original stratigraphic thickness is unknown due to poor exposure and complex internal structure, but it has an apparent composite structural thickness of at least 2.5 km. The Big Hill shale is typical of the flysch facies rocks that form the vast majority of the Izee terrane. Two tuff layers within the sandy flysch of the Big Hill shale near Mineral, Idaho, have yielded zircon ages of 180.61 ± 0.17 and 173.91 ± 0.07 Ma (Tumpane, 2010). These ages are consistent with normal stratigraphic order across the boundary between the upper Huntington and Weatherby formations and suggest a depositional relationship between the units. The presence of tuffs in the lower Weatherby Formation also suggests that Olds Ferry volcanism may have persisted into the early Middle Jurassic. Stops during Day 1 of the field trip will provide opportunities to examine directly various components of the Olds Ferry terrane, including the volcanosedimentary rocks of the lower Huntington Formation (Stop 1.2), intrusive rocks of the Brownlee pluton that form the oldest known component of the terrane (Stop 1.3), and the contact zone between the upper Huntington Formation and Weatherby Formation of the Izee terrane near the Bay Horse Mine (Stop 1.4). Stop 1.5 will examine a typical example of the flysch that makes up the majority of the Weatherby Formation, and Stop 1.6 will look at argillite and limestone of the Baker terrane as well as the Connor Creek Thrust that forms the contact between the Izee and Baker terranes. GEOLOGY OF THE WALLOWA TERRANE The constituent rocks of the Wallowa terrane range in age from the Permian to Early Cretaceous (Vallier, 1995). Exposures of crystalline basement rocks of the Wallowa arc occur in several locations within the Blue Mountains Province, especially within the Snake River Canyon from Oxbow, Oregon, downstream to its confluence with the Salmon and Imnaha rivers. Stratified silicic volcanic and volcaniclastic sequences of the Permian (Guadalupian) Windy Ridge and Hunsaker Creek formations comprise the oldest supracrustal rocks in the Wallowa terrane (Vallier, 1967, 1977, 1995). These rocks appear to be unconformably overlain by thick middle to upper Triassic (Ladinian and Carnian) volcanic and volcaniclastic rocks of mafic to intermediate composition known as the Wild Sheep Creek and Doyle Creek formations (Vallier, 1967, 1977, 1995). Late Triassic (earliest Norian) and Early Jurassic massive carbonate and sandstone-mudstone flysch sequences unconformably overly the Permian and Triassic volcanic assemblages (Vallier, 1977, 1995). This field trip will examine the geology of the Wallowa terrane near Pittsburg Landing, Idaho, where midcrustal arc rocks in the Cougar Creek Complex, have been structurally juxtaposed

Blue Mountains Province, Oregon and Idaho with fluvial to marine rocks of Jurassic age (White and Vallier, 1994; Schmidt et al., 2009). Recent U-Pb zircon geochronology of igneous rocks in the Wallowa terrane provides a useful framework of age information (Table 2), and together with geochemical and isotopic data, help to delineate the timing and petrologic characteristics of magmatism in this terrane (Vallier, 1995; Tumpane, 2010; Kurz, 2010). In the following section, we provide an overview of field relationships, geochronology, and geochemistry of rocks in the Wallowa terrane, with a particular focus on plutonic rocks of the Cougar Creek Complex and the Jurassic volcanosedimentary section at Pittsburg Landing. Cougar Creek Complex The Cougar Creek Complex is one of several major basement complexes located within the Blue Mountains Province and is exposed along ten kilometers of the Snake River in Hells Canyon from Temperance Creek northward to Pittsburg Landing (Fig. 4). Vallier (1995) described exposures similar to the Cougar Creek Complex in the Salmon River Canyon, near White Bird, Idaho, and again further to the northeast along the South Fork of the Clearwater River near Grangeville, Idaho. Assuming these northeastern exposures are continuations of the Cougar Creek Complex, the complex represents a regionally significant tectonic feature (Vallier, 1995). The Cougar Creek Complex is petrologically diverse, containing dikes and small plutons with a range of bulk compositions, including gabbro, basalt, diorite, quartz diorite, tonalite, and their metamorphosed and deformed equivalents (Vallier, 1995; Kurz, 2001; Kurz and Northrup, 2008). Metamorphism and deformation occurred under lower to upper greenschist facies conditions (Avé Lallemant, 1995; Vallier, 1995; Kurz, 2001; Kurz and Northrup, 2008; Kohn and Northrup, 2009). Principal tectonic fabrics include foliation and mylonitic shear zones that strike

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northeast-southwest, and dip moderately to steeply to the northwest and southeast directions. Stretching lineations are subhorizontal to gently plunging to the southwest and northeast (White and Vallier, 1994; Vallier, 1995; Avé Lallemant, 1995; Kurz and Northrup, 2008). Ductile shear zones that range from one centimeter to several meters in width are dominantly sinistral with minor conjugate right-lateral senses of shear, and were generated within a transpressional tectonic environment characterized by predominantly strike-slip to oblique-slip kinematic conditions (Avé Lallemant et al., 1985; Avé Lallemant, 1995; Kurz, 2001; Kurz and Northrup, 2008). The northern margin of the Cougar Creek Complex is formed by the Klopton Creek fault, which places plutonic rocks of the complex structurally over volcaniclastic and shallow marine rocks of the Coon Hollow Formation (Fig. 4; White et al., 1992; White and Vallier, 1994; Vallier, 1995; Kurz, 2001; Schmidt et al., 2009). The Klopton Creek fault is interpreted as a rightlateral oblique thrust fault and has been grouped with two similar faults in the region, namely the Hammer Creek fault, located to the east in the Salmon River canyon near White Bird, Idaho, and the Mount Idaho fault located to the southeast of Grangeville, Idaho (Schmidt and Lewis, 2007; Schmidt et al., 2007; Garwood et al., 2008; Kauffman et al., 2008). Together, these faults constitute a single, large-scale, north- to northeast-trending structure referred to as the Klopton Creek–Hammer Creek–Mount Idaho fault zone (Schmidt and Lewis, 2007). This fault system offsets the western Idaho shear zone but is stitched by the Bitterroot lobe of the Idaho Batholith. Thus, it has a maximum age of ca. 90 Ma, and a minimum age of 55–60 Ma (McClelland and Oldow, 2007; Schmidt and Lewis, 2007; Giorgis et al., 2008). The Trudy Mountain gneissose unit forms the most northern part of the Cougar Creek Complex (Fig. 4). Layering reflects intrusive relationships between older silicic rocks and younger, more mafic rocks, forming intricate dikes and screens of material

TABLE 2. RECENT U-Pb ZIRCON GEOCHRONOLOGY OF IGNEOUS ROCKS, WALLOWA TERRANE 206 238 Sample Geologic Rock Pb/ U Age identification unit type (Ma) Cougar Creek Complex and related plutonic rocks CC-7-17-1 Klopton Creek pluton CC08-05 Trudy Mountain gneissose unit CC07-04 Suicide Point pluton CC07-02 Suicide Point pluton CC08-01 CC-8-3-1 OX08-02 OX08-08 CC08-06 CC08-03 SAL09-01

Trudy Mountain gneissose unit Triangle Mountain pluton Wildhorse quartz diorite Oxbow quartz diorite Trudy Mountain gneissose unit Trudy Mountain gneissose unit Salmon River Canyon pluton

Gabbro Diorite Quartz diorite dike Gabbro

229.13 ± 0.45 229.37 ± 0.08 229.43 ± 0.08 229.43 ± 0.08

Protomylonitic tonalite Tonalite Quartz diorite Quartz diorite Mylonitic tonalite Mylonitic tonalite Tonalite

248.75 ± 0.08 254.21 ± 0.14 258.74 ± 0.06 259.14 ± 0.18 262.69 ± 0.10 262.80 ± 0.17 268.57 ± 0.07

Pittsburg Landing volcanosedimentary series ≤159.62 ± 0.10 07MB05 Conglomerate and sandstone unit Hbl-phyric lapilli tuff 07BM06 Upper red tuff unit Welded tuff 196.82 ± 0.06 Note: All dates are from U-Pb zircon thermal ionization mass spectrometry analyses. Data for the Cougar Creek Complex are from Kurz (2010), and data for the Pittsburg Landing series are from Tumpane (2010). See these sources for detailed sample locations, complete data tables, and analytical methods.

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116° 28' W

Lithologic Units Quaternary Undifferentiated Material Miocene Columbia River Basalt Group Late Jurassic Coon Hollow Formation Late Triassic Wild Sheep Creek Formation

44 32 16

45° 38' N

Klopton Creek Fault

Late Triassic Klopton Creek Pluton

Snake River

32 82

Late Triassic Kirby Creek Unit

U

U

19 57

U

K

t on op

38

U

T

Cr e ek

Permian - Triassic (?) Granitic Fault Sliver

U

il

D

D

D

D

ra

Late Triassic Trudy Mountain Mafic Unit

D

k ee Cr

Upper Pittsburg Landing

l

Late Triassic Suicide Point Pluton

78 36

Late Permian Triangle Mountain Pluton

85 72

79

Middle Permian - Late Triassic Kirkwood Creek Dike Unit Middle Permian - Late Triassic Trudy Mountain Gneissose Unit Permian Hunsaker Creek Formation 80

45° 36' N

Foliation

Stretching Lineation

50

61 29 U

re

e k

G

ar ug

k

o rg e

84 U

30

D

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83

74 D

Cr e

st V a lley C k

Kirkwood Fault

Lo

66

U

ood

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ul c h K irkw

U D

88

G

86 74

m a

40 16

29 35 34

86 70

c

Su

Gu

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l ch

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C

er

l

T e mp

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ht

C o rr

l

C re ek

66

72

72

58 U D

29 by

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Kirkwood Bar

68

75 72

ny

Triangle Mountain

G u lch

R o al y

C ree

69

69 79

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nce e re

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reek

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Kir

D

D

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ug

T wo

Sa

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H o mi

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ek

80

la

Two Corral Creek Fault U

59

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44

Ca

U

87 C

70

C

Co

88 t

D Lower Kirby Rapids

M

Magmatic Lineation (Mineral)

a

72

26

U

ek

Fault, Approx.

74

re

Cleavage

45° 34' N

rra

eek

u ir

Fault, Concealed

S

Cr

14

116° 30' W

Conctact

66 82

C re

Bedding

Thrust Fault

Du r ha m

6

Contact, Approx.

U D

68

Co

Lithologic Contacts and Structural Features

Fault

7

75

116° 26' W

Geochronology Sample

116° 32' W

74

ek

78

.

Suicide Point

25

88

U D

82

72 66

C r ee k

Study Area

24

73

83 76

52

75 Du

nc

22

an G

ul ch

80 12

h

45° 32' N

54 yG Dr

ulc

0

1

2 Kilometers

Figure 4. Geologic Map of the Cougar Creek Complex (from Kurz, 2010). Geochronologic data are summarized in Table 2.

Blue Mountains Province, Oregon and Idaho that have been variably deformed. Layers of both age groups range from a few centimeters to several meters in thickness and occasionally occur as large tabular bodies several tens of meters in thickness. Essentially all intrusive bodies within this unit are oriented parallel to the dominant northeast-southwest structural trend of the Cougar Creek Complex. Because compositional banding resulted from magmatic interlayering rather than metamorphic differentiation, we have applied the term “gneissose” to this unit. The southern boundary of the gneissose unit is gradational with the Trudy Mountain mafic unit, which is defined by a greater abundance of mafic material, a lack of relatively older screens of silicic rock, and a generally weaker deformation. The Trudy Mountain mafic unit includes gabbro, diorite, diabase, and minor quartz diorite dikes. Fine-grained basalt dikes are ubiquitous, cross-cutting most other intrusive units within the Cougar Creek Complex. Although the Trudy Mountain mafic unit is partly distinguished from the gneissose unit by an overall decrease in the amount of strain, similar deformational features are present. The southern boundary between the mafic unit and the Kirby Creek unit is also gradational, characterized by a transition from more mafic lithologies to predominance of quartz diorite dikes. Farther to the south, the mafic unit is structurally juxtaposed with adjacent units along the Two Corral Creek fault, a right-lateral oblique reverse fault likely related to the Klopton Creek fault. New high-precision U-Pb zircon geochronology of plutonic rocks from the Cougar Creek Complex and other basement exposures were generated by Kurz (2010) and are summarized in Table 2. These ages constrain the timing and duration of magmatism in the Wallowa arc terrane. Wallowa plutonic rocks record two compositionally and temporally distinct episodes of magmatic activity. From the Middle Permian to the Early Triassic (265.35 ± 0.18 Ma to 248.75 ± 0.08 Ma), the Wallowa arc was dominated by silicic calc-alkaline magmatism. The time interval from Early to Late Triassic represents an apparent gap in magmatic activity. In the Late Triassic (229.43 ± 0.08 Ma to 229.13 ± 0.45 Ma), magmatism was renewed briefly, and it was dominated by mafic to intermediate tholeiitic compositions. Geochemical and trace element characteristics imply derivation of the Late Triassic magmas from a previously depleted source unlike that of the older silicic intrusive bodies. Intrusive rocks associated with the older Middle Permian to Early Triassic cycle of magmatism follow a calc-alkaline fractionation trend, and trace element concentrations normalized to normal mid-oceanicridge basalt (MORB) show typical depletions in high field strength elements, indicating a subduction zone setting (Kurz, 2010). In contrast, intrusive rocks associated with the Late Triassic (ca. 229 Ma) cycle of magmatism are mafic and have tholeiitic geochemical affinity. Sr, Nd, and Pb isotopic data for intrusive rocks from both cycles of magmatism in the Wallowa arc show consistently juvenile compositions, in contrast with the isotopically more evolved rocks from the Olds Ferry arc (Kurz, 2010). Episodes of intrusive magmatism in the Cougar Creek Complex and other basement complexes correspond temporally and

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compositionally with volcanic components of the supracrustal stratigraphy of the Wallowa terrane (Vallier, 1967, 1977). The older Middle Permian to Early Triassic (265.35 ± 0.18 Ma to 248.75 ± 0.08 Ma) episode of silicic calc-alkaline activity corresponds, in part, with the known Guadalupian fossil age of the Hunsaker Creek Formation (Vallier, 1967, 1977). Given the generally similar compositions and timing, these supracrustal and basement rocks may be equivalent and record the same Middle Permian to Early Triassic cycle of magmatism. Based on available fossil data, the Hunsaker Creek Formation may not extend into the Early Triassic (Vallier, 1967, 1977). However, the top of the formation is truncated by an unconformity which separates it from the overlying Late Triassic Wild Sheep Creek Formation (Vallier, 1967, 1977). Thus, part of the original volcanic series equivalent to the Early Triassic intrusive rocks of the Cougar Creek Complex may have been removed by erosion prior to deposition of overlying units. This apparent unconformity has been described and correlated throughout the Wallowa terrane (Vallier, 1977, 1995), and the depositional hiatus coincides with the apparent cessation of magmatism in plutonic rocks of the Wallowa terrane from 248.75 ± 0.08 Ma to 229.43 ± 0.08 Ma. The Late Triassic (ca. 229 Ma) cycle of intrusive activity correlates, in part, with the latest Middle to early Late Triassic (Ladinian to Carnian) fossil ages of the Wild Sheep Creek Formation of the Seven Devils Group (Vallier, 1967, 1977). The Wild Sheep Creek Formation is overlain by and interfingered with the Late Triassic Doyle Creek Formation, and these units have been interpreted to represent a single interval of volcanic deposition (Vallier, 1977). Unfortunately, as with the Hunsaker Creek Formation, precise age constraints for the total duration of volcanic activity represented by the Wild Sheep Creek and Doyle Creek Formations are not known. The Doyle Creek Formation is overlain by the Martin Bridge Formation of earliest Norian age (Vallier, 1977; Follo, 1994; LaMaskin et al., 2008a). New paleontological investigations of the Martin Bridge Formation show that deposition of this thick carbonate began in the latest Carnian (e.g., Stanley et al., 2009), and by inference volcanic activity associated with the Wild Sheep Creek and Doyle Creek Formations had slowed or ceased entirely prior to this time (Vallier, 1977, 1995; Follo, 1992; Stanley et al., 2009). Jurassic Section, Pittsburg Landing The Pittsburg Landing area in western Idaho and eastern Oregon contains rocks of Permian to Miocene age (Fig. 5; White and Vallier, 1994; Kurz, 2001; Kurz et al., 2009; Schmidt et al., 2009). Complete descriptions of these units can be found in Vallier (1977), with updates to unit boundaries and names in White and Vallier (1994). The exposed volcanic and sedimentary section includes the Big Canyon Creek unit of the Wild Sheep Creek Formation, the Kurry unit of the Doyle Creek Formation, and the Coon Hollow Formation. The Big Canyon Creek unit at Pittsburg Landing is dominated by basalt and basaltic andesite pillow lavas, hyaloclastic sediments, and pillow breccias but also

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contains coarse-grained volcaniclastic rocks and massive flows, all of which are intercalated with conglomerate, sandstone, mudstone, and tuff beds (White and Vallier, 1994; Schmidt et al., 2009). Although the Doyle Creek Formation regionally consists of andesitic and rhyolitic lava flows as well as pyroclastic and epiclastic deposits, the Kurry unit of this formation in the Pittsburg Landing area consists largely of tuffaceous sandstone and mudstone with smaller amounts of volcanic breccia, sandstone channel-fill deposits, argillaceous limestone, tuff, and conglomerate (White and Vallier, 1994; LaMaskin et al., 2008a; Schmidt et al., 2009). In other parts of the Wallowa terrane, the Late Triassic (Carnian-Norian) Martin Bridge Limestone and Late Triassic to Early Jurassic (Norian-Toarcian) Hurwal Formation overlie

the Doyle Creek Formation, but these units are absent at Pittsburg Landing (Follo, 1992, 1994; Stanley et al., 2009; Vallier, 1977). The Coon Hollow Formation overlies the volcanic rocks of the Seven Devils Group (Big Canyon Creek and Kurry units) along an angular unconformity (Vallier, 1977; White and Vallier, 1994; Schmidt et al., 2009). The Coon Hollow Formation, as traditionally interpreted at Pittsburg Landing, consists of four main units: a basal red tuff unit, a fluvial conglomerate and sandstone unit, a marine sandstone and mudstone unit, and a turbidite unit, as well as some small intrusive bodies (Fig. 5). New U-Pb zircon geochronology from the Pittsburg Landing area provides the first radiometric age controls on the timing of deposition of the Coon Hollow Formation, and suggests

Figure 5. Generalized geologic map of the Pittsburg Landing area (after Tumpane, 2010; modified from White and Vallier, 1994). The locations of samples 07BM06, at the top of the red tuff unit, and 07BM05, near the base of the conglomerate and sandstone unit, are shown, and geochronologic data are summarized in Table 2. The conglomerate and sandstone unit of White and Vallier (1994) is interpreted as the base of a fluvial-deltaic to marine transgressive series represented by the Coon Hollow Formation. For full descriptions of rock units, see White and Vallier (1994). A more detailed geologic map of this area is given by Schmidt et al. (2009).

Blue Mountains Province, Oregon and Idaho revision of the stratigraphic nomenclature. The Kurry unit of the Doyle Creek Formation underlies the red tuff unit below an angular unconformity. Age control for the Kurry unit is based on ammonite molds and Halobia fossils that give an early Carnian to early Norian age, ca. 235–215 Ma (?) (White and Vallier, 1994; LaMaskin et al., 2008a). The red tuff unit consists of ~35 m of conglomerate and sandstone overlain by ~15 m of tuff beds. This unit shows evidence of erosion and cut and fill structures with coarse clastic material in scoured channels. The welded nature of the tuff in the upper part of the unit indicates subaerial deposition (Fig. 6). U-Pb zircon geochronology of a 2-m-thick layer of welded tuff located immediately below the angular unconformity that marks the base of the overlying cobble and sandstone unit yields a weighted mean 206Pb/238U age of 196.82 ± 0.06 Ma (Tumpane, 2010). This date places the age of the red tuff unit in the Early Jurassic (Sinemurian), significantly older than the previously accepted Middle Jurassic (Bajocian) age indicated by White and Vallier (1994). Because the red tuff unit is significantly older than the Coon Hollow Formation and separated from it by an angular unconformity, the red tuff

Figure 6. Field photographs of 1.5-m-thick welded tuff in the upper portion of the Red Tuff unit. (A) shows general outcrop character; (B) shows detailed view of outcrop surface with visible phenocrysts and flattened pumice fragments.

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unit must be regarded as distinct from the Coon Hollow Formation in the stratigraphic section. The red tuff unit was deposited sometime after the early Carnian to early Norian Kurry unit (ca. 215 Ma; LaMaskin et al., 2008a) and its depositional interval included 196.82 Ma (early Sinemurian). The conglomerate and sandstone unit of White and Vallier (1994) was deposited in angular unconformity on the red tuff unit, and we regard it as the base of a fluvial-deltaic to marine transgressive sequence represented by the Coon Hollow Formation. This unit contains ~1–15-m-thick conglomeratic beds with sub- to well-rounded, well-sorted clasts of volcanic (lava flow), volcaniclastic, and plutonic origin. The conglomerate beds are thickest in paleochannels and grade laterally and vertically into sandstone and mudstone with abundant plant debris, and the abundance of conglomerate beds decreases stratigraphically higher in the unit (White and Vallier, 1994; Schmidt et al., 2009). Several meter-thick reworked crystal tuffs and tuffaceous sands

Figure 7. Field photographs of ash layers in the lower portion of the conglomerate and sandstone unit of the Coon Hollow Formation. (A) Brown sandstone in the lower part of the photo is overlain by ~40 cm of light colored ash, which is overlain by dark, heterolithic pebble conglomerate. (B) Sharp basal contact of ash deposited on dark, fine sandstone with soft-sediment loading structures. Grain size in the ash layer fines upward.

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occur in cross-bedded channels and thin laterally into normally graded tuff horizons (Fig. 7). U-Pb zircon geochronology of a hornblende-phyric lithic lapilli tuff located ~40 m above the base of the conglomerate and sandstone unit yields a maximum depositional age of 159.62 ± 0.10 Ma (Tumpane, 2010). It could not have been deposited prior to the Oxfordian stage of the Late Jurassic. Thus, the conglomerate and sandstone unit of the lower Coon Hollow Formation is significantly younger than the previous Bajocian age estimates (White and Vallier, 1994; LaMaskin et al., 2008a). Oxfordian Coon Hollow strata were previously thought to occur only at the type locality, at Little Cougar Creek and Coon Hollow (Follo, 1992; Vallier, 1977; White and Vallier, 1994; White et al., 1992), and their presence at Pittsburg Landing strengthens the correlation of stratigraphy here with the type section. Within the conglomerate and sandstone unit, the coarsegrained deposits appear to be event beds interspersed with sandstone layers. Both of these components fine up section and transition conformably into the marine sandstone and mudstone unit of White and Vallier (1994). The marine sandstone and mudstone unit is dominated by sandstone near its base and changes to mudstone higher in the unit, but it is found only in limited outcrop at Pittsburg Landing (Schmidt et al., 2009). Shallow-water coral, pelecypod and brachiopod fossils in this unit are the basis for the Bajocian age assigned to the Coon Hollow section (Follo, 1992; Imlay, 1986; Stanley and Beauvais, 1990; White and Vallier, 1994). The turbidite unit occurs only as a fault-bounded package on the Oregon side of the Snake River and does not have any readily identifiable stratigraphic relationships to the other units of the Coon Hollow Formation, although it has been interpreted as a deeper water equivalent of the sandstone and mudstone unit (White and Vallier, 1994; Schmidt et al., 2009).

the early Late Triassic, from ca. 230 Ma to 210 Ma, a second cycle of magmatism occurred, consisting of mafic to silicic, tholeiitic to calc-alkaline activity that became isotopically more evolved with time. This phase is represented by the lower Huntington Formation and equivalent plutonitic rocks. After the emplacement of the Iron Mountain granodiorite (ca. 210 Ma), and until the deposition of the upper member of the Huntington Formation, the Olds Ferry terrane records a second episode of erosion and apparent hiatus in igneous activity. The initiation of explosive silicic volcanism associated with the upper Huntington Formation (Brooks, 1979b; Vallier, 1995; Tumpane, 2010) represents the third episode of igneous activity in the Olds Ferry arc, which began sometime after 210 Ma and lasted at least until 187 Ma, and perhaps until 174 Ma (early Middle Jurassic; Tumpane, 2010). Magmatism in the Wallowa Arc The plutonic rocks of the Wallowa arc record two temporally and compositionally distinct episodes of magmatic activity (Kurz, 2010). From the Middle Permian to the Early Triassic (ca. 265 Ma to 248 Ma), the Wallowa arc produced silicic calc-alkaline magmatism. From the Early to Late Triassic, an apparent hiatus in magmatic activity exists in the lithologic record of the arc. In the Late Triassic (ca. 229 Ma), brief but voluminous magmatism was dominated by mafic to intermediate tholeiitic compositions. The Middle Permian to Early Triassic and Late Triassic episodes of magmatism in the Wallowa arc correspond, both compositionally and temporally, with the Hunsaker Creek and Wild Sheep Creek Formations of the Seven Devils Group (Vallier, 1967, 1977), respectively. The Wild Sheep Creek Formation overlies the Hunsaker Creek Formation above an apparent unconformity (Vallier, 1977, 1995; Mann and Vallier, 2007), coinciding with the Early Triassic to Late Triassic hiatus in magmatism.

DISCUSSION Comparison of Olds Ferry and Wallowa Arcs New field observations, high-precision U-Pb zircon geochronology, trace element geochemistry, and tracer isotope analyses provide a more complete framework in which to evaluate the timing, duration, and tectonic context of magmatism in the Olds Ferry and Wallowa arc terranes. In this section, we describe, compare, and contrast the magmatic histories of the Olds Ferry and Wallowa arc terranes. Magmatism in the Olds Ferry Arc At least three episodes of magmatism and two unconformities occur in the geologic record of the Olds Ferry arc from the Middle Triassic to the Early Jurassic. First cycle magmatic activity consisted primarily of mafic to silicic tholeiitic magmatism of Middle Triassic age, and is represented by the Brownlee pluton (ca. 237 Ma) and other, relatively old plutonic clasts contained in conglomerates and breccias of the Lower Huntington Formation. From ca. 237 Ma to no later than ca. 230 Ma, the Olds Ferry arc underwent a period of uplift and erosion with no apparent record of igneous activity. In

Middle Permian to Early Triassic magmatism is widespread in the Wallowa, but not identified in the Olds Ferry. Magmatism in both arcs occurred in the Late Triassic, but with contrasting geochemical signatures and durations. Sr, Nd, and Pb isotopic data for magmatic rocks of the Olds Ferry arc terrane consistently display more enriched compositions compared to intrusive units of the Wallowa arc. Thus, we interpret the genesis of magmatic rocks of the Olds Ferry arc to have involved some degree of interaction with enriched continental material, either via subducted sediments and subsequent modification of the mantle wedge, and/or through assimilation of continental crust during magma emplacement. In contrast to the Olds Ferry arc, igneous rocks of the Wallowa arc are consistently depleted, indicating derivation from a depleted MORB mantle source within an intra-oceanic setting. Isotopic data of igneous rocks in the arc terranes from Kurz (2010) and Tumpane (2010) are consistent with those of LaMaskin et al. (2008b) and Schwartz et al. (2010), which were derived largely from sedimentary rocks in the Olds Ferry, Wallowa, and Baker terranes.

Blue Mountains Province, Oregon and Idaho Terrane Correlations Links between constituent terranes of the Blue Mountains Province and Paleozoic to Mesozoic lithotectonic elements along the western North American Cordillera have been difficult due to a lack of geochronologic, geochemical, and isotopic constraints. In the past, the Wallowa, Baker, and Olds Ferry terranes have been correlated with accreted terranes to the north as a collective tectonostratigraphic assemblage. For example, the Wallowa, Baker, and Olds Ferry terranes comprise a west to east triad consisting of magmatic arc–argillite-matrix mélange–magmatic arc, respectively, similar to those documented in the Intermontane superterrane of the Canadian Cordillera (e.g., from west to east: the Stikinia arc terrane–Cache Creek argillite-matrix mélange terrane–Quesnellia arc terrane). Lithologic, faunal, structural, and temporal similarities between these northern arc-related terranes and those in the Blue Mountains Province allow a compelling north to south continuation of the group (Mortimer, 1986; Stanley and Senowbari-Daryan, 1986). The Stikinia and Quesnellia arc terranes have been variably interpreted as a single magmatic belt (Church, 1975; Monger, 1977; Monger and Church, 1977; Dostal et al., 1999, 2009), or as separate volcanic arcs (Mortimer, 1986). These interpretations have led to a range of tectonic models that attempt to explain their present day positions in relation to one another and relative to the intervening Cache Creek mélange. Among these are (1) the right-lateral offset of a single arc (Wernicke and Klepacki, 1988; Beck, 1991, 1992; Irving et al., 1996), (2) the oroclinal closure of a single arc and consumption of the intervening ocean basin (Nelson and Mihalynuk, 1993; Mihalynuk et al., 1994), (3) Middle Jurassic thrusting of the Cache Creek terrane over the arc and subsequent synclinal folding (Samson et al., 1991; Gehrels et al., 1991), (4) the extrusion of high-pressure rocks of the Cache Creek terrane into the central portion of the arc (Dostal et al., 2009), and (5) the collision of two distinct arcs, an outboard Stikinia arc and an inboard Quesnellia arc (Mortimer, 1986). Sr, Nd, and Pb isotopic data from the Blue Mountains province support a peri-cratonic arc setting for the Olds Ferry arc (Dickinson, 1979; Miller, 1987; Dorsey and LaMaskin, 2007; Schwartz et al., 2010), similar to interpretations for the Stikinia and Quesnellia arcs (Dostal et al., 2009, and references therein). In addition, magmatic rocks from the Olds Ferry are isotopically similar to Late Triassic extrusive and intrusive rocks from both Stikinia and Quesnellia (Ghosh, 1995; Smith et al., 1995; Dostal et al., 1999, 2009). The lower Huntington Formation is lithologically and temporally similar to the Middle to Upper Triassic stratigraphy of Stikinia and Quesnellia arcs, including arc-related, mafic to intermediate volcaniclastic units, massive flows, and epiclastic rocks deposited in subaerial and submarine environments (i.e., the Takla, Stuhini, and Nicola Groups; Mortimer, 1987; Monger et al., 1991; Ferri and Melville, 1994; Pantaleyev et al., 1996; Dostal et al., 1999; MacIntyre et al., 2001; Beatty et al., 2006; Dostal et al., 2009). Middle to Late Triassic vol-

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canogenic assemblages from Stikinia, Quesnellia, and the Olds Ferry (i.e., the lower Huntington Formation) are also bounded by similar aged unconformities (Dostal et al., 2009). Stratigraphic, lithologic, and temporal similarities between Late Triassic assemblages from these different arc terranes provide reasonable evidence for their association, and representation as a single continuous fringing arc system in the Late Triassic and possibly into the Early Jurassic. Provided the interpretation that Stikinia and Quesnellia represent portions of the same arc, and following a north to south correlation between the Wallowa–Baker–Olds Ferry and Stikinia–Cache Creek–Quesnellia groups, requires that the Wallowa and the Olds Ferry terranes are also portions of a single arc. However, U-Pb ages, geochemical data, and isotopic data from the Wallowa and Olds Ferry arcs show that temporally overlapping Late Triassic magmatism from these arcs are fundamentally different (Kurz, 2010). During this time, a brief but voluminous episode of dominantly thoeliitic magmatism in the Wallowa arc occurred contemporaneously with generally calc-alkaline magmatism in the Olds Ferry arc (Kurz, 2010). Furthermore, Sr, Nd, and Pb isotopic data clearly distinguish the Middle Permian to Late Triassic intra-oceanic Wallowa arc from an isotopically enriched fringing Olds Ferry arc (Kurz 2010). This geochemical and isotopic distinction precludes any model that employs a single Wallowa–Olds Ferry arc system. Exploring possible scenarios of separate Wallowa and Olds Ferry arcs leads to previous interpretations of the Wallowa arc representing a fragment of Wrangellia, exposed in the Vancouver Islands of British Columbia and in the Wrangell and St. Elias Mountains of southeastern Alaska (Jones et al., 1977; Hillhouse et al., 1982; Wernicke and Klepacki, 1988, Dickinson, 2004). Kurz (2010) described similar lithostratigraphic characteristics between Wrangellia and the Wallowa, such as Late Paleozoic arc and marine sequences that are subsequently overlain by voluminous, predominantly mafic volcanic sequences above a regional unconformity. Middle to Late Triassic basalt sequences from both the Wallowa arc terrane (e.g., Wild Sheep Creek Formation) and Wrangellia (e.g., Nikolai and Karmutsen Formations) are also nearly synchronous, and were erupted over a relatively short interval of time (Kurz, 2010). Normal MORB and chondritenormalized trace element patterns for extrusive rocks of the Wallowa and Wrangellia are similar, and may be linked through the common mechanism of spreading ridge subduction (Kurz, 2010). Following this interpretation, the Wallowa represents an outboard intra-oceanic arc distinct from the Olds Ferry. FIELD GUIDE Day 1. Geology of the Olds Ferry Terrane Stops examined on the first day provide an overview of the geology of the Olds Ferry terrane. Driving directions. The field trip begins at Farewell Bend State Park, Oregon. To access the park, take Interstate 84 to exit

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353 and drive north on State Highway 30 for ~1 mi. The entrance to Farewell Bend State Park is located on the east side of the road. Stop 1.1. Farewell Bend State Park 44.3047 N, 117.2265 W Farewell Bend is located at the southern end of the Snake River Canyon, which cuts northward through the Blue Mountains region and provides relatively continuous exposures of the Paleozoic-Mesozoic basement. The canyon transects obliquely across the composite arc terranes and allows relationships between the first-order tectonic units to be examined. After leaving Farewell Bend State Park, the field trip will travel on relatively remote roads for ~40 mi before reaching Richland, Oregon. Driving directions. Continue north on State Highway 30 for ~4.2 mi to Huntington, Oregon. Turn right onto Washington Street, which becomes the Snake River Road (SRR) Proceed eastward on the SRR following the Burnt River toward Brownlee reservoir. Beginning ~0.4 mi after the SRR crosses to the north side of the Burnt River, fairly continuous exposures of the Huntington Formation can be found on the hillslope and roadcuts north of the road. Stop 1.2 is located along the SRR, from ~1.0 to 1.35 mi past the crossing of the Burnt River. Stop 1.2. Lower Huntington Formation, Confluence of the Burnt River and Snake River 44.3628 N, 117.2383 W The lower member of the Huntington Formation is well exposed east of the town of Huntington near the confluence of the Burnt and Snake Rivers (Fig. 3). Exposures of the Old Ferry terrane continue northward along the Snake River for ~11 km. Massive lava flows and interflow breccias predominate, with minor volcanic sandstones and pyroclastic rocks. Massive lavas in the lower Huntington Formation are typically porphyritic andesite to basaltic andesite with large plagioclase phenocrysts and glomerocrysts (Fig. 8). Hydrothermal alteration and low-grade metamorphism produced abundant secondary chlorite and epidote. Tuff breccias are common, and some include dacite and rhyodacite clasts in addition to the ubiquitous andesite and basaltic andesite. Limestone pods and lenses are present sporadically in the lower Huntington Formation. Many are rounded or mechanically broken and may be olistostromal; however, some form laterally persistent beds and may represent localized primary deposition of bioclastic carbonate within the section (LaMaskin, 2008; Tumpane, 2010). The Huntington Formation as a whole commonly has been assigned an age of Late Triassic (Carnian-Norian) based on fossil data (Brooks, 1979a, 1979b; Vallier, 1995; Dorsey and LaMaskin, 2007; LaMaskin, 2008). Recent recalibrations of the Triassic time scale (Ogg et al., 2008; Walker and Geissman, 2009), suggest that ammonites of the lower Huntington have an age of ca. 232 Ma. New direct dating of volcanic rocks in the Huntington Formation by U-Pb zircon geochronology, however, indicates that these biostratigraphic data may provide a maximum age for deposition of the lower Huntington and do not reflect the complete age range of

the formation. U-Pb zircon geochronology of two volcanic rocks in the Lower Huntington Formation yielded ages of 221.72 ± 0.12 and 220.66 ± 0.18 Ma (Table 1; Tumpane 2010). Driving directions. Continue driving east and then north on the SRR for ~3.0 mi to the Brownlee pluton. Stop 1.3. Brownlee Pluton 44.3965 N, 117.2247 W At this stop, we will examine the Brownlee pluton and its contact relationships with the adjacent Huntington Formation. The Brownlee pluton is a relatively small body of trondhjemite outcropping on both the west and east sides of the reservoir (Figs. 9 and 10). The pluton has been interpreted as intrusive into the Huntington Formation (Juras, 1973; Brooks et al., 1976; San Filippo, 2006) or as basement upon which the Huntington Formation was deposited (Vallier, 1995; Dorsey and LaMaskin, 2007). Recent detailed mapping of the Brownlee pluton and surrounding area has helped to clarify field relationships with surrounding units and gives new insight into the lithostratigraphic architecture of the Olds Ferry terrane (Tumpane, 2010; Kurz, 2010). Contact relationships are consistent with the presence of two unconformities: one at the contact between the Brownlee pluton and lower Huntington Formation, and a second at the base of the upper Huntington Formation. Fine-grained basaltic to basaltic andesite dikes and sills cross-cut the trondhjemite pluton and represent feeder dikes to Huntington volcanic rocks (Fig. 10). Key outcrops are located at 480977 E, 4916464 N in Zone 11T, NAD27 UTM coordinates (Tumpane, 2010). Reddish oxidation exists in the Brownlee trondhjemite below the contact with overlying lava flows, and brecciation with incipient soil development occur at the top surface of the pluton. Several meters above this contact, volcanic rocks of the lower Huntington Formation are truncated by brown, moderately

Figure 8. Basaltic andesite tuff breccia typical of the lower Huntington Formation, showing numerous mm- to cm-size phenocrysts of plagioclase.

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Figure 9. Geologic map of the Brownlee Pluton and surrounding area after Kurz (2010), compiled from Tumpane (2010), Brooks (1979a), and Juras (1973).

Figure 10. Field photo showing typical outcrops of the Brownlee pluton with m-scale cross-cutting mafic dikes. View looks northeast at outcrops on the east side of the Snake River.

sorted volcaniclastic deposits of the upper Huntington Formation. The erosional surface at the base of the upper Huntington cuts obliquely across the section, placing the upper Huntington in direct depositional contact with both the lower Huntington and the Brownlee pluton along strike (Fig. 9). The radiometric age of the Brownlee pluton (237.68 ± 0.07 Ma; Kurz, 2010) is older than the oldest known components of the lower Huntington Formation, consistent with the interpretation of the Brownlee pluton as depositional basement for the overlying volcano-sedimentary sequence. The maximum duration of the unconformity at the base of the lower Huntington Formation is ~6 m.y., based on the difference in age between the Brownlee pluton and the late Carnian age (ca. 232 Ma; LaMaskin, 2008) of ammonites in overlying limestone pods of the lower Huntington Formation. The duration of the unconformity separating the lower and upper members of Huntington Formation is not well constrained, because the precise age of the youngest

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parts of the lower member and the depositional base of the upper member are not known. Driving directions. From the Brownlee pluton, continue north on the SRR ~4.4 mi to the Bay Horse mine. Along this traverse, the road cuts obliquely up section through the lower and upper Huntington Formation. The mine is located on the hill slope ~100 m east of the SRR. To access the mine area, park along the SSR and walk ~400 m up the small access road that cuts sharply back to the left (east). Stop 1.4. Bay Horse Mine 44.4541 N, 117.2188 W The Bay Horse mine area contains exposures of the upper Huntington Formation and the basal section of the overlying Weatherby Formation of the Izee terrane (Figs. 11 and 12; Livingston, 1925; Henricksen, 1975). The series of rocks observed at this location is characteristic regionally of the upper Huntington Formation. The stratigraphically lowest rocks locally exposed are a package of volcanic, volcaniclastic, and sedimentary rocks, including a coarse volcanic breccia and lava flow

Figure 11. Field photograph of turbidites interstratified with more massively bedded volcanic rocks of the upper Huntington Formation.

overlain by an ~2-m-thick layer of siltstone, shale, and sandstone. Above this package are a pair of volcanic units: ~10 m of rhyodacite breccias and interbedded lava flows, and 15 m of rhyolite tuff. Clast size in the rhyodacite breccias is graded, decreasing from up to 50 cm at the base of the unit to 1–3 cm at the top. The interiors of the thin (~1 m thick) lava flows are plagioclase phyric and vesicular, and flow bases are lobate, consistent with deposition in a submarine environment. The rhyolite tuff produces a prominent cliff-forming layer at the level of the mine entrance. The volcanosedimentary rocks below the rhyodacite are typical of the upper Huntington Formation, with 1–10-m-scale packages of coarse, poorly sorted volcanic breccia interbedded with decimeter-scale, well-bedded sediments and occasional lava flows. The rhyodacite and rhyolite tuff are interpreted as conformable with the underlying rocks and therefore, represent the top of the upper Huntington sequence at this location. These felsic volcanic rocks can be correlated with thicker sequences of similar composition and tectonostratigraphic position located ~20 km to the northeast in the Dennett Creek area of Idaho, where the rhyodacite and rhyolite have yielded U-Pb zircon dates of 188.45 ± 0.05 Ma and 187.03 ± 0.04 Ma, respectively (Tumpane, 2010) Orientations of bedding in the upper Huntington Formation and the overlying McChord Butte conglomerate of the Weatherby Formation indicate an angular discordance of ~30° between these units across their contact near the Bay Horse mine. The McChord Butte conglomerate contains locally derived clasts of the underlying rhyolite tuff and rhyodacite, as well as other volcanic, plutonic, and sedimentary rocks (Brooks, 1967, 1979a, 1979b; Henricksen, 1975; Juras, 1973; Mann and Vallier, 2007). Driving directions. From the entrance road to the Bay Horse mine, continue northward on the SRR ~7.1 mi to Stop 1.5, located

Figure 12. Field photograph looking north from the Bay Horse mine showing the contact between the upper Huntington and basal Weatherby formations (from Tumpane, 2010).

Blue Mountains Province, Oregon and Idaho where the SSR jogs abruptly westward as it contours through the entrance to a large side canyon. The Weatherby Formation is well exposed in roadcuts.

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Turn north on U.S. Highway 95 and continue ~80 mi to Riggins, Idaho. Overnight in Riggins. Day 2. Geology of the Wallowa Terrane

Stop 1.5. Flysch of the Weatherby Formation 44.5293 N, 117.1817 W This stop provides an opportunity to examine the flysch facies of the Weatherby Formation, by far the most volumetrically significant component of this unit. Penetrative cleavage is ubiquitous in the shale and mesoscopic rootless folds suggest the possibility of complex internal structure. However, the slope-forming nature of the Weatherby shales makes an analysis of its internal structure difficult. The age of the lower Weatherby Formation is constrained by the presence of a tuff located ~50 m above the depositional base of the unit dated at 180.61 ± 0.17 Ma; however, the age of youngest deposition in the Weatherby Formation is not known, and its upper contact with the overlying Baker terrane is structural (Stop 1.6). Driving directions. Continue northward on the SRR approximately 6.6 mi to Stop 1.6, located where strongly deformed argillite and limestone are well exposed in a roadcut. Stop 1.6. Connor Creek Thrust—The Boundary between the Izee and Baker Terranes 44.5969 N, 117.1304 W Outcrops and roadcuts at this stop contain complexly deformed argillite, phyllite, and limestone characteristic of the Baker terrane. Several generations of deformational fabric are evident, including penetrative foliation, crenulation cleavage, and folds in various orientations. To the east (across the reservoir), a prominent limestone layer with complex internal structure dips steeply to the north. This unit is near the structural base of the Baker terrane. The Connor Creek thrust sits ~100 m below the limestone and places the chert argillite, phyllite, greenstone, and limestone of the Baker terrane structurally over the mudstones and shale of the Weatherby Formation. The Connor Creek thrust, therefore, marks the boundary between the Baker and Izee terranes at this location. The marine lithologies and complex structure of rocks in the Baker terrane are consistent with assembly of this terrane in a forearc accretionary prism environment. Due to the fact that the Baker terrane lies between the Wallowa and Olds Ferry arc terranes, the paleogeographic association of these forearc components with one or both volcanic arcs is difficult to reconstruct. Based on trace-element and Sr and Nd isotope geochemistry, Schwartz et al. (2010) inferred that much of the Baker terrane accretionary complex formed in association with the Olds Ferry arc in a pericontinental setting. Driving directions. Continue northward on the SRR ~17 mi to Richland, Oregon. Turn right and head east on State Road 86 for ~28.4 mi to Oxbow and Copperfield. Turn right and travel south on the Brownlee-Oxbow highway for ~11 mi. Cross the Snake River just below Brownlee dam, and continue on Idaho State Highway 71 south and east to Cambridge, ID, ~29 mi.

Field trip stops on Day 2 will focus on the geology of the Wallowa terrane in the vicinity of Pittsburg Landing, Idaho. This area contains midcrustal plutonic rocks of the Cougar Creek Complex (Stop 2.2) as well as fluvial to marine rocks of Jurassic age (Stop 2.3). These units are juxtaposed structurally by the Klopton Creek fault, which will be discussed in more detail during an introduction and overview of the Pittsburg Landing geology (Stop 2.1). Driving directions. From Riggins, Idaho, drive northeast on U.S. Highway 95 for ~27 mi. Just before the highway bends to the right to leave the river and climb the Whitebird grade, turn left onto River Road and continue for ~1 mi to the bridge over the Salmon River. Cross the bridge and turn left (south) onto Deer Creek Road/National Forest Development Road 493. Follow the Deer Creek Road/ National Forest Development Road 493 for ~17 mi to Pittsburg Landing. Take the minor access road southward to a parking lot located near the hiking trail head to Kirby Creek (45.6228 N, 116.4643W). Walk a few hundred meters to the southeast, climbing to the wind gap at the top of the nearby small hill. Stop 2.1. Pittsburg Landing—Overview of Wallowa Terrane 45.6210 N, 116.4597 W The top of the hill provides an excellent vantage point to view and discuss the geology of the Pittsburg Landing area. To the south and east, massive outcrops of Permian and Triassic plutonic rocks of the Cougar Creek Complex can be seen. To the north and west are well-stratified rocks of Jurassic age. The abrupt topographic break separating these domains follows the NE-SW–trending trace of the Klopton Creek fault, a regionally significant structure of probable Cretaceous age (White and Vallier, 1994; Schmidt et al., 2009). The fault thrusts midcrustal rocks of the Cougar Creek Complex over the Jurassic section. The rest of the day will be spent exploring the geology of the Cougar Creek Complex (Stop 2.2) and the Jurassic section at Pittsburg Landing (Stop 2.3). Directions. Follow the hiking trail along the east side of the Snake River southward from Pittsburg Landing into the Cougar Creek complex. Stop 2.2. Cougar Creek Complex 45.6133 N, 116.4625 W The transition into the complex is clearly expressed by a change in the morphology of the canyon. The crystalline basement rocks of the Cougar Creek Complex are competent, outcrop well, and produce a narrower, steep-sided canyon. The Cougar Creek Complex is petrologically diverse, containing dikes and small plutons with a range of bulk compositions, including gabbro, basalt, diorite, quartz diorite, and tonalite (Vallier, 1995;

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Kurz, 2001; Kurz and Northrup, 2008). After crossing the Klopton Creek fault and entering the gorge, outcrops of the Trudy Mountain gneissose unit are common. Layering is subvertical and reflects intrusive relationships between older silicic rocks and younger, more mafic rocks, forming intricate, anastomosing series of dikes and screens that have been variably deformed

(Fig. 13). The lighter colored silicic screens represent the Middle Permian to Early Triassic (ca. 268–248 Ma) phase of calcalkaline magmatism found in the Wallowa terrane, and the mafic dikes were products of the Late Triassic (ca. 229 Ma) pulse of dominantly mafic tholeiitic magmatism (Kurz, 2010). Deformation is heterogeneously developed, and commonly concentrated in anastomosing mylonitic shear zones ranging from a few mm to cm in thickness. The shear zones are subvertical and subparallel to the compositional layering of the gneissose unit. Stretching lineations are nearly horizontal, and kinematic indicators viewed in outcrop surfaces and oriented thin sections suggest a dominantly left-lateral sense of structural transport (Kurz, 2001; Kurz and Northrup, 2008). Left-lateral mylonitic shearing affected Middle Permian to Early Triassic tonalitic rocks, as well as Late Triassic mafic to intermediate intrusive bodies. This deformation is more intensely recorded in older tonalitic rocks relative to intrusive units of the younger mafic cycle of magmatism. Discrete zones of deformation in Late Triassic bodies often exhibit strain gradients, and are cross-cut by other weakly deformed mafic to intermediate units. Field observations and geochronology of syndeformational titanite indicate that this deformational event overlaps with the Late Triassic (ca. 229 Ma) cycle of magmatism (Kurz, 2010; Kurz and Northrup, 2008). Directions. Return to Pittsburg landing via the hiking trail. From the trail head parking lot (45.6228 N, 116.4643 W), either drive or walk north on the access road 1.2 mi to the intersection with the main gravel road leading to the campground. Turn right and travel northeast ~0.25 mi to where a small dirt road branches to the left (northwest). Park at the junction of these roads. Stop 2.3. Jurassic Section, Pittsburg Landing 45.6228 N, 116.4643 W This stop is a walking traverse through part of the Jurassic section and will examine the red tuff unit and the conglomerate and sandstone unit of White and Vallier (1994). Walk down the road to the northwest for ~0.6 mi. The hill to the northeast has outcrops of the red tuff unit on the lower part of the slope, overlain by lighter colored stratigraphy of the cobble and sandstone unit (Fig. 14). Beginning from where the small road makes a sharp lefthand bend (45.6471N 116.4735W), climb the hill to the northeast, working progressively through the section. The red tuff unit consists of ~35 m of conglomerate and sandstone overlain by ~15 m of tuff beds. This unit shows evidence of erosion and cut and fill structures with coarse clastic material in scoured channels, and the welded character of the tuff layers in the upper portion of the

Figure 13. Field photographs of the Trudy Mountain gneissose unit showing: (A) subvertical sheets and screens of variably deformed tonalite (ca. 270–250 Ma, Permian–Early Triassic) separated by more weakly deformed mafic dikes (ca. 229 Ma, Late Triassic), and (B) detailed relationship between non-deformed gabbro and mylonitic tonalite (sample CC08-03).

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streamline and focus the text. Thanks to Erika Akin for providing insightful comments, careful proofreading, and good editorial feedback at various stages in the development of this manuscript. Much appreciation to Tracy Vallier for sharing his insights, knowledge, and enthusiasm for the geology of the Blue Mountains arc terranes. REFERENCES CITED

Figure 14. Field photograph looking northeast showing the angular unconformity between the red tuff unit and overlying conglomerate and sandstone unit of the Coon Hollow Formation (after Tumpane, 2010; photo by Reed Lewis).

unit is consistent with subaerial deposition. Welded tuff located immediately below the angular unconformity that marks the base of the overlying cobble and sandstone unit has been dated at 196.82 ± 0.06 Ma (Fig. 6; Tumpane 2010). The conglomerate and sandstone unit of White and Vallier (1994) was deposited in angular unconformity on the red tuff unit, and contains ~1–15-m-thick conglomeratic beds with sub- to well-rounded, well-sorted clasts of volcanic (lava flow), volcaniclastic, and plutonic origin. Abundance of conglomerate beds decreases stratigraphically higher in the unit. Several meterthick reworked crystal tuffs and tuffaceous sands occur in crossbedded channels and these thin laterally into normally graded tuff horizons (Fig. 7). U-Pb zircon geochronology was conducted on one of the tuff horizons located ~40 m above the base of the conglomerate and sandstone unit, yielding a maximum depositional age of 159.62 ± 0.10 Ma (Tumpane, 2010). Continue to climb the small ridge line, moving obliquely across the stratigraphy of the cobble and sandstone unit, reaching a location near the top of a small drainage on the southeastern side of the hill (45.6484N 116.4672W). Outcrops in this drainage show spectacular examples of the volcanic ash horizons in the lower part of the cobble and sandstone unit. Meter-scale cross stratification can be seen in thicker layers. Thinner ash layers show grading and soft sediment loading structures. Walk carefully down through the section along this drainage and return to the vehicles. ACKNOWLEDGMENTS Jeff Lee and Todd LaMaskin provided careful, complete, and helpful reviews of this field guide, which helped greatly to

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Proceedings of the Circum-Pacific terrane conference: Stanford, California, Stanford University Publications, Geological Sciences, p. 197–199. Wagner, N.C., Brooks, H.C., and Imlay, R.W., 1963, Marine Jurassic exposures in Juniper Mountain area of eastern Oregon: Bulletin of the American Association of Petroleum Geologists, v. 47, p. 687–701. Walker, J.D., and Geissman, J.W., compilers, 2009, Geologic Time Scale: Geological Society of America, doi: 10.1130/2009.CTS004R2C. Wernicke, B.P., and Klepacki, D.W., 1988, Escape hypothesis for the Stikine Block: Geology, v. 16, p. 461–464, doi:10.1130/0091-7613 (1988)016<0461:EHFTSB>2.3.CO;2. White, J.D.L., White, D.L., Vallier, T.L., Stanley, G.D., Jr., and Ash, S.R., 1992, Middle Jurassic strata link Wallowa, Olds Ferry, and Izee terranes in the accreted Blue Mountains island arc, northeastern Oregon: Geology, v. 20, p. 729–732, doi:10.1130/0091-7613(1992)020<0729:MJSLWO> 2.3.CO;2. White, D.L., and Vallier, T.L., 1994, Geologic evolution of the Pittsburg Landing area, Snake River canyon, Oregon and Idaho, in Vallier, T.L., and Brooks, H.C., eds., Geology of the Blue Mountains region of Oregon, Idaho, and Washington—Stratigraphy, physiography, and mineral resources of the Blue Mountains region: U.S. Geological Survey Professional Paper 1439, p. 75–89. Wilson, D., and Cox, A., 1980, Paleomagnetic evidence for tectonic rotation of Jurassic plutons in Blue Mountains, eastern Oregon: Journal of Geophysical Research, v. 85, p. 3681–3689, doi:10.1029/JB085iB07p03681.

MANUSCRIPT ACCEPTED BY THE SOCIETY 9 MARCH 2011

Printed in the USA

The Geological Society of America Field Guide 21 2011

Neogene drainage development of Marsh and Portneuf valleys, eastern Idaho Glenn D. Thackray David W. Rodgers Andrew Drabick Department of Geosciences, Mail Stop 8072, Idaho State University, Pocatello, Idaho 83209, USA

ABSTRACT Neogene drainage development in southeastern Idaho has been influenced by drainage capture, Basin and Range faulting, volcanism, and the Late Pleistocene Lake Bonneville overflow and Bonneville Flood. In Marsh Valley, the Middle to Late Pleistocene sedimentary sequence is dominated by alternating lacustrine/paludal and alluvial sediments, which have yielded new 40Ar/39Ar, amino acid racemization, and luminescence age estimates. The pattern of sedimentation through time indicates poor drainage integration of southern Marsh Valley through most of the last ca. 640 ka and suggests slow basin subsidence along Quaternary faults mapped on the basin edges. Marsh Valley initially incised into that valley fill sequence ca. 19 ka, shortly before the Bonneville Flood. Marsh Creek is a markedly underfit stream occupying a meandering, broad valley carved into the valley fill sequence. These geomorphic and sedimentologic patterns suggest non-catastrophic Lake Bonneville overflow before and after the Bonneville Flood. In Portneuf Valley, ca. 8.5–7.4 Ma basin fill and a bedrock pediment are perched 800 m above the modern valley floor. Major incision of basin fill and bedrock by the ancestral Portneuf drainage system occurred prior to the Middle to Late Pleistocene, when two cut-fill events resulted in accumulation of alluvial fan deposits extending ~10–60 m above the modern valley floor and basalt extending ~10 m below to 20 m above the modern valley floor. Final incision by Lake Bonneville overflow is evident but relatively minor in comparison to the cumulative downcutting. Overall, incision is attributed to isostatic subsidence of the eastern Snake River Plain, which served as base level for the Portneuf drainage system after passage of the Yellowstone hot spot in late Miocene time.

Thackray, G.D., Rodgers, D.W., and Drabick, A., 2011, Neogene drainage development of Marsh and Portneuf valleys, eastern Idaho in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 89–101, doi:10.1130/2011.0021(04). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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INTRODUCTION Neogene basins of southeast Idaho formed within the active Basin and Range province and on the shoulders of the initially tumescent and then subsiding Yellowstone–Snake River Plain magmatic center. The drainage systems in the region thus reveal a complex history of drainage capture, active faulting, and volcanism, capped by the latest Pleistocene Bonneville Flood. Here we describe the basin evolution of Marsh Valley and Portneuf Valley, two major basins in the Portneuf River drainage that exemplify that complex history. The field guide traces the path of Marsh Creek and the Portneuf River, which also represents the path of the Bonneville Flood. The guide examines evidence for base level change, drainage capture, and the Bonneville Flood. GEOLOGIC SETTING Several Cenozoic basins mark the intersection of the northeastern Basin and Range province with the eastern Rocky Mountain Region and the Snake River Plain. Marsh and Portneuf Valleys in southeast Idaho are examples of these Cenozoic basins (Fig. 1). The major physiographic features in both these valleys include dissected horsts of Proterozoic and Paleozoic bedrock, broad flat-bottomed valleys, and transverse valleys (e.g., Portneuf Gap) that transect mountain ranges and connect valleys (Ore, 1982). Middle and Late Pleistocene geologic features including lava flows, overflow features from the Bonneville Flood, and extensive valley-bounding benches with linear edges. Together, these features reveal the recent geologic history of this region. In southeast Idaho, the late Cenozoic structural basins are characterized by a multi-stage history outlined by Ore (1982). The initial stage is the tectonic formation of the basins by Basin and Range faulting during late Miocene time. The valleys are principally half-grabens and the corresponding ranges are horsts of late Neoproterozoic and Paleozoic marine sediments. These valleys were initially filled with detritus from uplifted horst blocks as well as volcanic ash derived from rhyolite eruptions on the Snake River Plain. Initial excavation of this basin fill reflects basin capture and base level changes, which Rodgers et al. (2002) attributed to subsidence of the Snake River Plain. A final incision and filling episode, associated with Lake Bonneville overflow before, during, and after the Bonneville Flood, created the present valley floor levels in both Marsh and Portneuf Valleys. Marsh and Portneuf Valleys are structural basins bounded by ranges consisting largely of Neoproterozoic and Paleozoic sedimentary rocks (Figs. 1 and 2). North-striking, west-dipping normal faults cut the bedrock to define half-grabens, with the faults showing several kilometers of cumulative slip. The resulting grabens filled with the late Miocene and Pliocene Salt Lake Group comprising variable amounts of fluvial, lacustrine, and volcaniclastic sediments (Oriel and Platt, 1980; Ore, 1982; Link et al., 1985). In Marsh Valley, the Middle to Late Pleistocene formation of Marsh Valley (Long and Link, 2007) overlies the Salt Lake

Group and is composed of fluvial, alluvial, and lacustrine sediments (Shields, 1978; Gaylen, 1978; and Steele, 1980) that are locally difficult to discern from the Salt Lake Group (Kruger et al., 2003). In Portneuf Valley the Pleistocene fluvial, alluvial, and lacustrine sediments are subhorizontal and more poorly consolidated than the east-tilted, well-indurated Salt Lake Group. Gravity studies in both Marsh and Portneuf valleys detail the structure and sub-surface morphology. Henrich (1979) and Cumming (1980) concluded that Marsh Valley is a full graben structure with a depth of 2.5 km in the south and 1.7 km depth in the north. Kruger et al. (2003) combined geophysical investigations with geologic mapping from Crane (2000), Pope et al. (2001), and Janecke and Evans (1999) to produce a generalized structural map of the Marsh Valley area showing that the valley is an asymmetric full graben with a major normal fault on the east side and a minor normal fault on the west side of the valley. The most recent geophysical study of Portneuf Valley (Reid, 1997) showed that the valley is divided into two basins separated by a gravity high at Red Hill. That gravity high corresponds to a segment boundary in the range-bounding Fort Hall Canyon fault (Rodgers and Othberg, 1999). Investigation of the Quaternary history of Marsh Valley originated with G.K. Gilbert’s (1890) initial investigation into the development and subsequent overflow of Lake Bonneville. While Gilbert apparently did not enter Marsh Valley, he did infer that an outburst flood had occurred. O’Connor (1993) and Malde (1968) subsequently described the path of the Bonneville Flood, characterized its deposits, and inferred its hydraulics. Several masters’ degree students and other authors have completed work in both Portneuf and Marsh Valley areas (Gaylen, 1978; Shields, 1978; Steele, 1980; Pope, 2002; Bush, 1980; Reid, 1997; Kruger et al., 2003). Most of this work has focused on geophysical research into valley structure, with limited focus on Quaternary deposits. Shields (1978) and Steele (1980), however, examined small outcrops of the formation of Marsh Valley and inferred depositional environments based on the material exposed. Despite the long history of research in Marsh Valley, several aspects of the Quaternary development of Marsh Valley remained unknown until recently, including timing of depositional and erosional events, correlation of stratigraphic units on a broad scale, factors controlling the depth of the basin, relation of the sedimentary and geomorphic record to fluctuations of Pleistocene Lake Bonneville, and the relationship and sensitivity of the valley to Snake River Plain tectonics, to local volcanism, and to climatic fluctuations. These questions were addressed in the yetto-be-finished Idaho State University M.S. thesis work of Andrew Drabick, on which much of this field guide contribution is based. Drabick conducted field work in 2005–2006, with assistance from Glenn Thackray and David Rodgers. Geochronologic data were produced by Darrell Kaufman at Northern Arizona University (amino acid racemization on shells, AAR), Tammy Rittenour at Utah State University and Ron Goble at University of Nebraska (optically stimulated luminescence, OSL), and Brent Turrin at Rutgers University (40Ar/39Ar dating of rhyolitic tephras).

Neogene drainage development of Marsh and Portneuf valleys PLUVIAL LAKES AND BONNEVILLE FLOOD Basin and Range faulting, coupled with altered climate patterns during Quaternary glaciations led to the formation of numerous pluvial lakes in closed basins throughout the province. Pleistocene Lake Bonneville was the largest of these pluvial lakes and at its largest extent covered 51,530 km2 in western Utah and Parts of Nevada and Idaho (O’Connor, 1993). Latest Pleistocene lake level rise in the Bonneville Basin is attributed

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to two distinct events: the first is the diversion of the Bear River around 50 ± 10 ka (Bouchard et al., 1998), which increased discharge into the basin by 33% and the second is cooler, moisture conditions during the Late Pleistocene (Bright, 1963; McCoy, 1987; Bouchard et al., 1998). Lake level rose until it reached its maximum elevation of 1552 m near Red Rock Pass. The lake level was maintained at this elevation for at least several hundred years until the flood was initiated at ca. 17 ka (14.5 14C ka; Malde, 1968; O’Connor, 1993; Benson et al., 2011). We speculate

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below that non-catastrophic overflow may have initiated several thousand years earlier. At Red Rock Pass, the lowest point along the Lake Bonneville margin, alluvial fans resulting from erosion of the nearby Portneuf and Bannock Ranges joined above Tertiary aged tuffaceous sediments and Cambrian Limestone (Link et al., 1999).

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Figure 2. Generalized geologic map of a portion of SE Idaho showing major faults. Note the normal faults on the western margin of southern Marsh Valley, making that basin a full graben amongst more typical half-graben basins in the region Modified from Janecke and Evans (1999), Crane (2000), Kruger et al. (2003), and Eversaul (2004).

Neogene drainage development of Marsh and Portneuf valleys the presence of the Cambrian limestone (O’Connor 1993). When flow of the water shifted west into less consolidated sediments, the channel widened rapidly and caused slope failures in numerous locations, allowing for the breakout of the flood (Malde, 1968; O’Connor, 1993; Janecke and Oaks, 2007). The path of the flood originated at Red Rock Pass and traveled north into Marsh Valley and eventually through the Portneuf Gap and the Portneuf Valley before entering the Snake River Plain at Michaud Flats (Malde, 1968). The flood followed the path of the Snake River westward until meeting with the Columbia River and eventually into the Pacific Ocean. Evidence of the path of the flood beyond Lewiston is obscured by the channeled scablands associated with the Missoula Floods (O’Connor, 1993). The maximum flood discharges associated with the Bonneville Flood were between 0.85 and 1.15 million m3/sec at Red Rock pass and 0.57 to 0.62 million m3/sec near Lewiston (O’Connor, 1993). Variations in stream power were dictated by existing topography along the flood route. In narrow portions of the flood path, like the Portneuf Gap, flow was constricted and stream power increased by several orders of magnitude. This increased stream power removed large portions of the basalt of Portneuf Valley. Markedly lower stream power in areas like Marsh and Portneuf Valleys, where topography was wider and flatter, caused deposition of large volumes of gravel. NEW FINDINGS CONCERNING THE MIDDLE AND LATE PLEISTOCENE HISTORY OF MARSH VALLEY As noted, recent investigations have revealed new details of the Middle and Late Pleistocene stratigraphy of southern Marsh Valley and developed a chronology for the deposits. Here we highlight major findings of that investigation. Northern Marsh Valley contrasts sharply with southern Marsh Valley (Fig. 1). While southern Marsh Valley is broad and exposes thick accumulations of unconsolidated sediment, the northern valley is considerably narrower and exposes minimal unconsolidated sediment. The northern portion is dominated by the basalt of Portneuf Valley, which erupted 430 ± 70 ka (40Ar/39Ar, Rodgers et al., 2006) in Gem Valley, flowed past Lava Hot Springs, and entered Marsh Valley near present-day McCammon (Scott et al., 1982). Upon entering the broad middle portion of Marsh Valley, the main mass of the basalt flow turned north to fill the floor of northern Marsh Valley, the Portneuf Gap, and southern Portneuf Valley. A smaller portion of the basalt of Portneuf Valley ponded southward into Marsh Valley, creating a sill for the small amounts of water draining north from southern Marsh Valley. Drainage was diverted both west and east of the flow, creating a classic inverted topography. Southern Marsh Valley is dominated topographically by prominent benches cut by an incised broad, flat-bottomed meandering valley. Alluvial fan/bajada surfaces mark both the east and west sides of the valley. The western bench exposes alternating alluvial gravel and loess, and appears to merge topographically with a broad bench occupying the center of the valley and expos-

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ing fluvial and lacustrine/paludal sediments. The bench on the east side appears to rise above the central bench and thus may be older than the western and central benches. Marsh Creek is a markedly underfit stream occupying a broad valley (1–2 km) cut into the western/central bench surface. That valley extends southeast to the Red Rock Pass area, and is marked by alluvial terraces inset between the upper bench surfaces and the valley floor, particularly in southwest Marsh Valley. The general pattern of this incised valley along the length of southern Marsh Valley is that of a very long and broad meander (wavelength ~12 km, amplitude ~7 km). The morphology of this incised valley suggests that a large river flowed through the valley and, consequently, that Lake Bonneville overflowed its sill in the Red Rock Pass area before and/or after the Bonneville Flood. That is, we infer that Lake Bonneville overflowed at its sill, before the flood (sill at Bonneville lake level) and/or after the flood (sill at Provo lake level). As the Bonneville Flood entered southern Marsh Valley, the sudden loss of stream power caused deposition of extensive gravel deposits, eroded from the Red Rock Pass area, mantling the entire central bench (O’Connor, 1993). The main portion of the catastrophic flow was thus in a depositional mode in this area rather than an incisional mode, and the incision is more likely to have occurred when sediment load was lower and incisional capacity higher, i.e., by non-catastrophic, but significant, Bonneville overflow. Non-catastrophic overflow prior to the Bonneville Flood has also been suggested by S. Janecke (2010, personal commun.) and was implied by Currey (1982), and ca. 2.5 ka (17 cal ka– 14.5 cal ka) of non-catastrophic overflow was documented for the post-flood Provo lake level by Godsey et al. (2005). Recent sediment core-based work in the western Bonneville basin by Benson et al. (2011) suggests non-catastrophic overflow for ca. 1.5 ka (18.5–17 ka) before the flood and for ca. 1.8 ka (17– 15.2 ka) after the flood. The central and southern benches expose a laterally extensive sequence of Middle and Late Pleistocene lacustrine/paludal and alluvial sediments (these sediments are incorporated into a single stratigraphic unit known as the formation of Marsh Valley), dating from 637 ± 3 ka (40Ar/39Ar tephra age, Stop 1.1) to 19.22 ± 1.40 ka (UNL1435, OSL age, Stop 1.2). The overall stratigraphic pattern is of laterally and vertically extensive lacustrine/ paludal sediments with thin (<1.5 m) interbeds of alluvial sand and gravel. The lacustrine/paludal sediments consist largely of green, white, and light brown, generally massive silty clay and clayey silt, with locally common gastropod and bivalve shells. Together, the sediment character and mollusk taxa suggest dominance of a shallow marsh environment much like that along Marsh Creek in southern Marsh Valley today. Marsh deposition was interspersed with occasional aggradation of alluvial fan sand and gravel into the marsh environment. Thus, southern Marsh Valley appears to have been poorly drained and poorly integrated into the Portneuf River drainage through most of its known Middle and Late Pleistocene history. The primary exception to this pattern is the inferred period of non-catastrophic and catastrophic

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Bonneville overflow, during which sufficient discharge entered the valley to drain readily into the Portneuf drainage. Why has southern Marsh Valley remained poorly drained through the Middle and Late Pleistocene? We propose two explanations. First, the basalt of Portneuf Valley clearly acted as a sill after ca. 430 ka for drainage from southern Marsh Valley, and likely is partly responsible for poor drainage after that time. However, we have found paludal sediments as old as 637 ± 3 ka (40Ar/39Ar ash age), indicating poor drainage prior to emplacement of the basalt of Portneuf Valley. Our second explanation is that Marsh Valley experienced tectonic subsidence throughout the Middle and Late Pleistocene. As noted in the geologic background at the beginning of this section, southern Marsh Valley has a thick sediment fill and is bounded on its western edge by a normal fault with inferred Quaternary motion (Pope et al., 2001). Offset on that fault would induce subsidence of the floor of Marsh Valley and maintain poor drainage integration with the Portneuf River. Incision of the broadly meandering, modern Marsh Creek valley into the main bench surface appears to have occurred shortly after 20 ka, and may reflect initial, non-catastrophic Lake Bonneville overflow. As described with Stop 1.2 below, alluvial fan gravel is interbedded with overbank and/or eolian silt in the top 3–4 m of sediment below the bench surface on the southern edge of Marsh Valley. The uppermost gravels are in thin (<50 cm), discontinuous lenses suggestive of shallow channels cut into the fan surface. An OSL age of 19.22 ± 1.40 ka (UNL1435) was determined for a silt bed associated with the uppermost gravel bed. Thus, gravel deposition on the fan surface appears to have ceased around or shortly after 20 ka. We infer that the cessation of fan gravel deposition occurred when the small streams incised into the fan surface, in response to the incision of the valley of Marsh Creek into the fan surface (see above). If the incision of Marsh Creek into the fan surface was a result of non-catastrophic Bonneville overflow, as discussed above, then these inferences suggest that Lake Bonneville reached a level near its gravel dam at Red Rock Pass as early as ca. 20 ka. At that time, water would either have begun leaking through the gravel dam or overflowed the dam, creating a river in Marsh Creek, large enough to cause incision of the fan surface to isolate the broad benches marking southern Marsh Valley. As noted, Benson et al. (2011) suggested that Lake Bonneville overflowed at sill level for ca. 1.5 ka before the Bonneville Flood, starting ca. 18.5 ka. The OSL age noted here and Benson et al.’s (2011) inference are broadly compatible. PORTNEUF VALLEY LATE CENOZOIC GEOLOGY AND LANDSCAPE EVOLUTION Portneuf Valley is a small basin located northwest of Marsh Valley along the southern margin of the eastern Snake River Plain (ESRP) (Fig. 1). Like Marsh Valley, it initially formed in response to late Miocene Basin-Range tectonism but Portneuf Valley was more strongly influenced by ESRP tectonism due to its marginal location. As explained below, the modern Portneuf

Valley is interpreted to be deeply incised into the original, late Miocene Portneuf Valley due to base level changes associated with ESRP surface subsidence. The Bannock and Pocatello Ranges were originally one coherent horst uplifted on the west-dipping Arbon Valley fault, but this horst was subsequently cut by several normal faults including the Fort Hall Canyon fault (Figs. 3, 4). The age of faulting is constrained by the 8.5 Ma to <7.4 Ma age of a 900-m-thick section of conglomerate and tuff, equivalent to the Salt Lake Group, that accumulated in the hanging wall of the Fort Hall Canyon fault (Rodgers and Othberg, 1999). Faults and rocks including the Salt Lake Group were tilted 25° east by the time faulting ceased sometime after 7.4 Ma, creating the original Portneuf Valley. After deposition of the Salt Lake Group ended, Portneuf Valley primarily experienced incision and dissection. Incision is suggested by two features perched 800 m above the modern Portneuf Valley that we interpret as the floor of the original valley. One feature is an expanse of tilted Salt Lake Group at Justice Park in the Bannock Range, ~20 km south of Pocatello (Figs. 3, 5). The other is a 5 km2 pediment beveled into tilted bedrock just south of Kinport Peak, ~7 km southwest of Pocatello (Figs. 3, 5). Both features occupy the graben side of the Fort Hall Canyon fault and both features currently lie at an elevation of ~2150 m. As cross-cutting relations indicate the pediment post-dates Basin-Range tectonism, and as no younger faults are present, the simplest interpretation is that ~800 m of incision occurred to lower the valley floor to its current elevation of ~1350 m (Rodgers et al., 2002). The latest Miocene to Early Pleistocene history of incision is virtually unknown since no rocks of this age are preserved in the valley. After significant material was removed from Portneuf Valley and the adjacent Bannock and Pocatello ranges, two distinctive bajadas, now preserved as the East and West Benches (Fig. 3), formed at elevations of 1600–1420 m. These fan deposits of inferred Middle Pleistocene age (Scott et al., 1982) are ~50 m thick and overlie an unconformity that dips gently inward toward the valley axis (Fig. 4). Ore (1982) interpreted these relations as a Pleistocene cut-fill event in response to downstream base level change and/or climate fluctuations. We agree, and according to our interpretation the age and elevation of Benches provides evidence that most of the 800 m of valley incision occurred prior to the Middle Pleistocene (Fig. 6). The 430 ± 70 ka (40Ar/39Ar, Rodgers et al., 2006) Portneuf basalt flowed north along an existing river channel (Trimble, 1976) through northern Marsh Valley and Portneuf Valley, and thus documents incision to ~1350 m by that time. If the Portneuf basalt is younger than the bench deposits, it records a second cutfill event in which the Middle Pleistocene(?) bajada was incised by an axial river whose channel was then filled by basalt (Fig. 6) (Ore, 1982). Post-basalt incision in Portneuf Valley occurred to the west of the basalt and created the modern Portneuf River floodplain. Whether this channel was excavated continuously after ca. 430 ka or primarily during the 17 ka Bonneville Flood has not been

Neogene drainage development of Marsh and Portneuf valleys

Figure 3. Geologic map of the Portneuf Valley and surrounding region. Geology is discussed in text, and is simplified from Trimble (1976), Rodgers and Othberg (1999), and Link and Stanford (1999). Lines indicate location of cross sections shown in Figures 4 and 5. ESRP—eastern Snake River Plain.

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Figure 4. Transverse geologic cross-section of the southern Portneuf Valley. See Figure 3 for profile location. Lower section shows details of subsurface geology including several Neoproterozoic and Cambrian formations disconformably overlain by Salt Lake Group (Tsu) and downfaulted against the Neoproterozoic Pocatello Formation (Zp) along the Fort Hall Canyon normal fault. Upper section emphasizes near-surface Quaternary geology including fan gravels of the East and West Benches (Qfg, Ql/Qfgw), Portneuf basalt (Qp), boulder gravel of the Bonneville Flood (Qbg), post-flood alluvium (Qal), and post-flood alluvial fans (Qfp). From Rodgers and Othberg (1999).

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Neogene drainage development of Marsh and Portneuf valleys

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Figure 5. Profile of the Portneuf River superimposed on a simplified NNWSSE cross-section of the Bannock Range. See Figure 3 for profile locations. The top of the Salt Lake Group (Tsu) and Proterozoic (Z) bedrock near Kinport Peak are interpreted to define a latest Miocene pediment surface equivalent to the floor of the ancestral Portneuf Valley, which graded northward to an ancestral volcanic plateau on the eastern Snake River Plain (ESRP). After the volcanic plateau migrated northeast to Yellowstone, eastern Snake River Plain subsidence led to crustal flexure and surface downwarping along the ESRP southern margin. Ancient streams that flowed from the Basin-Range to ESRP, like the Portneuf River, adjusted to lowering base level by incising as much as 800 m. During the latest stages of incision, Quaternary-Pliocene basalt (QPb) accumulated on the ESRP. Modified from Rodgers et al. (2002).

determined from relations within the valley. Certainly the flood had a major impact—the constriction at Portneuf Gap created a hydraulic jump that increased stream power and led to significant erosion downstream (O’Connor, 1993). In the waning flood stages, considerable deposition of coarse gravels occurred within the modern floodplain and especially northwest of Portneuf Valley on the ESRP (Fig. 3). Post-Bonneville landscape evolution in Portneuf Valley has been relatively limited. A landslide, now occupied by the Johnny Creek subdivision, displaced the southern West Bench and destroyed the fluvial scarp along the channel’s west side (Fig. 4) (Rodgers and Othberg, 1999). Elsewhere the Benches are cut by transverse drainages with small fans at their mouths (Fig. 3), evidence of re-grading to the new (lower) local base level of ~1360 m. The modern Portneuf River is an underfit stream within its broad floodplain.

Figure 6. Proposed incision history of Portneuf Valley. Solid lines show measured elevations and ages of pediment (beveled into bedrock near Kinport Peak), sediments, and basalt. Dashed lines show interpreted incision history between these control points. Note change in scale along horizontal axis. See text for explanation.

Taken together, these geologic and geomorphic relations provide evidence for deep and widespread incision of the original Portneuf Valley, with most incision occurring from ca. 7 to 1 Ma (Fig. 6). Ore (1982) suggested that downstream base level changes induced the Pleistocene cut-fill events, and Rodgers et al. (2002) proposed the first-order cause of base level change was Snake River Plain surface subsidence. According to this model, an ancient analogue to the Yellowstone plateau was located north of Portneuf Valley from 7 to 10 Ma as manifested by rhyolite vent locations (Trimble, 1976; Pierce and Morgan, 1992; Kellogg et al., 1994). Northeast passage of the volcanic plateau after ca. 7 Ma, to its current location at Yellowstone, resulted in surface subsidence of the ESRP due to loading and thermal contraction (Brott et al., 1981; Rodgers et al., 1990, 2002; McQuarrie and Rodgers, 1998). Beginning sometime shortly after 7 Ma, the ancestral Portneuf drainage system probably reversed from south-directed to north-directed, and headward erosion of those streams—driven by ESRP surface subsidence—ultimately created the integrated drainage network of today. Evidence of the relation between ESRP subsidence and incision is provided by map-scale features in Portneuf Valley. First, rocks along the southern ESRP margin experienced northward flexure along a ENE axis, such that traces of older west-dipping structures like the Arbon Valley and Fort Hall Canyon faults bend abruptly northeast (Fig. 3). Second, the modern Portneuf Valley is elongated NNW, oblique to the original north-trending half-graben but perpendicular to the ESRP flexural axis. Consequently, incision of the original Portneuf Valley was not restricted to the more easily eroded Miocene basin fill but instead through both basin fill and bedrock, transverse to the original structural trend (Fig. 3). Both features are interpreted to be responses of the ESRP margin to subsidence of the ESRP province proper. FIELD GUIDE Stops in this field guide are arranged from south to north, starting in southern Marsh Valley. (This is the first day of a larger field trip for this meeting. See Link et al., this volume).The road

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log begins on U.S. Highway 91 at the south end of the town of Downey. Turn southwest on Old Malad Highway. Old Malad Highway traverses the central bench of Marsh Valley. This surface is underlain by Middle Pleistocene paludal and alluvial sediments and is capped by a veneer (1–6 m) of Bonneville Flood gravel (O’Connor, 1993). Drive 2.2 mi southwest on Old Malad Highway and turn left on Cherry Creek Road. The road here has descended into the flat valley floor of Marsh Creek. This valley continues NW, then north, through a series of broad meanders. Marsh Creek is markedly underfit in this valley, suggesting substantial river flow through Marsh Valley during the Late Pleistocene. Drive 0.4 mi south, and turn left (east) on East Back Downata Road. Drive 0.3 mi east to small roadside exposure on south side of road.

thin gravel layers. These relationships suggest that deposition on the alluvial fan ceased shortly after 20 ka, as discussed in the text. Return to Old Malad Highway, turn right, and return 3.8 mi northeast to Highway 91 at Downey, turn left (north) on Highway 91. Drive 0.5 mi north on Highway 91 to Woodland Road; turn left. Drive 2.8 mi west on Woodland Road, atop the central bench of Marsh Valley. Immediately beyond the I-15 interchange (and before the road descends into the Marsh Creek valley), turn right on Tool Road. Drive 1.6 mi, north then west, on Tool Road (a.k.a. South Ray Road), passing piles of drying peat excavated from the floor of Marsh Creek valley. After 1.6 mi, and on curve of road to north, pull over at small farm track descending off edge of bench on left side of road. Park here and walk down short track to exposures on left side. Note that this exposure lies on private land, and permission should be obtained before entering.

Stop 1.1. Small Roadside Exposure on Back Downata Road in Southern Marsh Valley (UTM 12 T 405059 E, 4694971 N)

Stop 1.3. Central Marsh Valley Stratigraphy in Farm Track off Tool Road (a.k.a. South Ray Road) (UTM 12T 402181E 4699858N)

This exposure, ~0.3 mi east of intersection with Cherry Creek Road, contains alluvial gravel and sand with an ~2-m-thick interbed of green clay, silt, and clayey silt. The sediments are typical of the alternating sequences of lacustrine/paludal sediments with alluvial sediments characteristic of southern Marsh Valley, although in most exposures the lacustrine/paludal sediments dominate. The prominent white layer in the center of the outcrop is a rhyolitic ash dated by 40Ar/39Ar methods to 637 ± 3 ka, and inferred to be Lava Creek Ash from the Yellowstone Plateau. Its thickness here is likely amplified by concentration of tephra eroded from the surrounding landscape. Travel 0.3 mi west on East Back Downata Road to Cherry Creek Road. Turn right on Cherry Creek Road, drive 0.1 mi and turn left to continue west on East Woodland Road. Drive 0.6 mi west on East Woodland Road to Old Malad Highway; turn left. Old Malad Highway here rises onto the southern bench of Marsh Valley, underlain primarily by alluvial fan gravels. Drive 0.6 mi south on Old Malad Highway, and turn left on South Bloxham Road. Drive 0.1 mi to enter gravel pit immediately adjacent to South Bloxham Road.

This track is cut into the bluffs of the central bench of southern Marsh Valley and exposes stratigraphy typical of this portion of the valley. The incised valley of Marsh Creek lies below. Note the poorly drained, marshy area with peat excavations. The section at this location is dominated by fine-grained paludal/lacustrine sediment with a fine pebble interbed and a carapace of presumed Bonneville Flood gravel. The section consists of a basal clay-silt unit 1.5 m thick with shells of Flumnicola and other taxa, overlain by 0.3 m of sandy pebble gravel, which yields sparse fragmental bone fossils, itself overlain by 2 m of light-colored clay-silt and coarsening upward into fine-medium sand. This upper clay-silt unit also yielded shells of Flumnicola and other taxa. Flumnicola from the lower shell-bearing bed yield an amino acid racemization age estimate of 116 ka, and from the upper shell-bearing bed an estimate of 138 ka. These age estimates are stratigraphically reversed, so the specific age estimates are not reliable, but the age estimates do suggest a latest Middle Pleistocene age for these strata. Flumnicola from the top of a nearby section yielded an age estimate of 107 ka. That section also contains abundant vertebrate fossils, including degraded tusks of mastodon or mammoth and bones of other mammals typical of Middle to Late Pleistocene fauna of the region. This stop also has a revealing view of the incised Marsh Creek valley and the west bench beyond. Note the breadth of this valley, its marshy character and poor drainage, and the alluvial terraces on the far side. These terraces suggest an extended period of incision and deposition by a sizeable river. The west bench of Marsh Valley visible from this location is part of an extensive bajada lying at the foot of the Bannock Range. A silty deposit (loess?) capping the upper gravel in the alluvial fan surface yielded an OSL age of 78.0 ± 6.0 ka (UNL 1434). Retrace route 1.6 mi back to I-15 interchange at Woodland Road. Proceed north on I-15. The next stop is ~30 mi north of this location in Pocatello. The following road log reveals pertinent

Stop 1.2. South Bloxham Road Gravel Pit (UTM 12T 403697E, 4694206N) This pit, now converted to a shooting range in classic East Idaho style, exposes sediments typical of the upper portions of alluvial fans in southern Marsh Valley. The exposure consists of several lenses of gravel that pinch and swell along the quarried face, suggesting shallow, gravel-bedded stream channels. Gravel beds generally coarsen upward and contain interbeds of fine sand and silt. This pattern is repeated through the 3–5 m pit wall exposure. The top of the quarried face is marked by a loess/ overbank silt bed. This bed yielded an OSL age of 19.22 ± 1.40 ka (UNL1435). Elsewhere in the face, this horizon is overlain by

Neogene drainage development of Marsh and Portneuf valleys features along that route. I-15 traverses the broad, flat central terrace of Marsh Valley. Note the gravel pits being excavated into Bonneville Flood gravels at various locations. Milepost 36: ~5 mi north on I-15 from the entrance point at Woodland Road, and a short distance north of Exit 36, the broad, meandering valley of Marsh Creek approaches the interstate from the southwest. Milepost 40: Four miles north of Exit 36, the interstate passes Arimo. Large gravel pits at the southern end of Arimo have been excavated into Bonneville Flood gravel bars noted by O’Connor (1993). Milepost 43: The 430 ka Portneuf basalt can be seen to the west of the interstate. The Portneuf basalt originated in Gem Valley, ~20 mi NE of here, and flowed down the Portneuf River drainage through Lava Hot Springs and then west into Marsh Valley, entering the valley ~2 mi NE of here. The lava mainly flowed north toward Inkom and Pocatello, but ponded southward into Marsh Valley, forming a sill for southern Marsh Valley. Mileposts 47–54: The interstate is located on top of the 430 ka Portneuf basalt. The basalt was scoured by the Bonneville Flood with minimal deposition of gravel. Glimpses of the incised Marsh Creek floodplain can be seen to the west (left) while the Portneuf River is incised in a narrow, hidden valley to the east (right). Beyond the two incised floodplains are muted fluvial scarps cut across the broad bajadas. Milepost 55: Bedrock highs in the valley center forced a constriction in the basalt flow. As the interstate drops off the basalt and into the confluence of Marsh Creek and Portneuf River, the basalt, river channels, and interstate all make a broad 90° bend and head west directly through the Bannock Range. Milepost 58–59: Interstate rises to top of Portneuf basalt. Incised Portneuf River valley visible to south (left), with prominent fluvial scarp cut into bedrock on south side. Milepost 60: Immediately south of here is a well-defined alluvial fan, sourced from the south, that post-dates the Bonneville Flood. Milepost 61: Note the small quarry in Bonneville Flood gravels adjacent to the south side of the interstate. A mile ahead is the Portneuf Gap, a prominent constriction that forces the highway, railroad, natural gas pipeline, and river into close proximity. Milepost 62: Portneuf Gap. Scour marks 100 m above the valley floor, especially on the south side, mark the highest level of the Bonneville Flood. Basalt is missing here but reappears a quarter-mile ahead where the interstate once again rises to its top. At this point, the road crosses the Fort Hall Canyon fault, a late Miocene normal fault that separates the Bannock Range from Portneuf Valley. Milepost 63: The highway turns northwest and Portneuf Valley is visible in all directions. Leave the interstate at exit 67. Turn left at stop sign and drive 1.0 mi north on South 5th Avenue. Turn right on Barton Road and drive 0.2 mi to base of Red Hill. Park on left side of road at trail sign (UTM 12T 384030E 4745233N). Take rough path with switchbacks to crest of Red Hill and walk ~700 m north

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of vehicles along the crest to its highest point, where the Snake River Plain first appears on the north horizon. Stop 1.4. Red Hill This end of Red Hill is underlain by Neoproterozoic Camelback Mountain Quartzite that dips gently east. Splays of the Fort Hall Canyon normal fault bound Red Hill on west and east sides, so that it forms an intrabasin bedrock high. To the north, the eastern Snake River Plain is a low-relief basaltic plain whose minor hills are shield volcanoes. The basalt near Pocatello is ca. 2–4 Ma (Kellogg et al., 1994; Tauxe et al., 2004). Most of the basalt is covered by younger alluvium and loess in the nearby Snake River Valley, but ~40 km away the basalt is exposed. Three prominent rhyolite (or rhyolite-cored) Pleistocene buttes are visible—Big Southern Butte, Middle Butte, and East Butte. On clear days the Basin-Range mountains north of the ESRP are visible on the horizon, ~100 km away. The ESRP formed in the wake of the Yellowstone hotspot and records ~6 m.y. of dominantly basalt magmatism and isostatic subsidence. To the east, the main strand of the Fort Hall Canyon fault is located at the base of the Pocatello Range where the break in slope occurs. The fault dips ~25° west beneath the Portneuf Valley including Red Hill. Several fault splays are present at the latitude of Red Hill, defining a segment boundary between single fault strands and sub-basins to the north and south. Cumulative offset is ca. 5 km (Rodgers and Othberg, 1999). To the south, the Fort Hall Canyon fault crosses Portneuf Valley at the base of Portneuf Gap and extends southward to Justice Park, located on the horizon at the break in slope ~500 m below and west of the tallest peak, Scout Mountain. Immediately west of the fault, the low hills are underlain by conglomerate and minor tuff of the late Miocene Salt Lake Group. Note that the elevation of exposed Salt Lake Group increases toward Justice Park, which is at ~2150 m. To the west, bedrock in the Bannock Range is everywhere tilted east due to block tilting during normal faulting. On the horizon just south of Kinport Peak (with the radio towers) is a flat pediment surface beveled into tilted bedrock. This surface is the same elevation (~2150 m) as Starlight Formation at Justice Park and interpreted to record the floor of the ancestral Portneuf Valley before late Cenozoic incision. Also to the west, in the foreground, is a bedrock high that gravity and drilling indicate is the westward continuation of Red Hill (Reid, 1997). From the bedrock northwards is the West Bench, a flat surface graded toward the center of the valley. The West Bench and the mirror-image East Bench are underlain by Middle Pleistocene alluvial fan deposits, evidence that an episode of backfilling occurred in the Portneuf Valley. Presumably the entire valley was first eroded to a level ~50 m below the benches, then backfilled to the level of the benches. A sharp fluvial scarp associated with the Bonneville Flood terminates the West Bench.

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The ca. 430 ka Portneuf basalt crops out just a few hundred meters southwest of southern Red Hill. This is the northernmost outcrop of the 70 km long basalt flow—from here it dips slightly north, is concealed by younger sediment for a few kilometers, and then terminates beneath northern Pocatello. During the 17 ka Bonneville Flood, water depth was sufficient to flow over the top of the Portneuf basalt but not over the West and East Benches. The top of Red Hill was probably just above water level. Boulders 2–3 m in diameter were transported short distances before final deposition in downtown Pocatello and Chubbuck, to the northwest of here. From Red Hill northward is a continuous layer of flood gravels characterized by particle sizes that decrease northward onto the Snake River Plain (Trimble, 1976). How much incision was completed by the main Bonneville Flood, or even its possible non-catastrophic precursor, has not been documented for this location. See also Link and Phoenix (1996), which contains historical photographs from this location and a summary of the Bonneville Flood in southeast Idaho. The modern Portneuf River floodplain is between the Portneuf basalt and the West Bench. The Portneuf River has established a meandering pattern along the valley floor, but has neither eroded nor deposited much sediment. However, its tributary streams have incised the foothills and East and West Benches and formed small alluvial fans on the valley bottom, as seen from this location. Return to vehicles. Return to Interstate 15 interchange and turn onto the northbound entrance ramp. Drive 25 mi north on Interstate 15 to Blackfoot and take exit 93 at the U.S. Highway 26 interchange. Drive northwest on U.S. 26 across the eastern Snake River Plain to Arco, a distance of ~60 mi, to find accommodations for the night. ACKNOWLEDGMENTS The Marsh Valley research was funded in part by NASA Experimental Program to Stimulate Competitive Research (EPSCoR) Grant # FPK302-02. Ar-Ar dating of volcanic tephra deposits was completed by Brent Turrin at Rutgers University. Amino acid racemization analyses were completed by Jordon Bright and Darrell Kaufman at Northern Arizona University. Optically stimulated luminescence dating was completed by Ron Goble at the University of Nebraska, and Tammy Rittenour of Utah State University assisted with sampling. This field guide benefited from reviews by Jeff Lee, Jennifer Pierce, and Paul Link, and from discussions on features at Red Rock Pass with Susanne Janecke. REFERENCES CITED Benson, L.V., Lund, S.P., Smoot, J.P., Rhode, D.E., Spencer, R.J., Verosub, K.L., Louderback, L.A., Johnson, C.A., Rye, R.O., and Negrini, R.M., 2011, The rise and fall of Lake Bonneville between 45 and 10.5 ka: Quaternary International, v. 235, no. 1–2, p. 57–69, doi:10.1016/j .quaint.2010.12.014.

Bouchard, D.P., Kaufman, D.S., Hochberg, A., and Quade, J., 1998, Quaternary history of the Thatcher Basin, Idaho, reconstructed from the 87Sr/86Sr and amino acid composition of lacustrine fossils-implications for the diversion of the Bear River into the Bonneville basin: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 141, p. 95–114, doi:10.1016/S0031 -0182(98)00005-4. Bright, R.C., 1963, Pleistocene Lakes Thatcher and Bonneville, southeastern Idaho [Ph.D. dissertation]: Minneapolis, Minnesota, USA, University of Minnesota, 292 p. Brott, C.A., Blackwell, D.D., and Ziagos, J.P., 1981, Thermal and tectonic implications of heat flow in the eastern SRP, Idaho: Journal of Geophysical Research, v. 86, no. B12, p. 11,709–11,734, doi:10.1029/ JB086iB12p11709. Bush, R.R., 1980, Gravity survey of Tyhee area, Bannock County, Idaho [M.S. thesis]: Pocatello, Idaho, USA, Idaho State University, 37 p. Crane, T.C., 2000, Geologic mapping and gravity survey of the Lava Hot Springs Idaho 7.5 min. quadrangle: Evidence for a Late Miocene supradetachment basin in southeast Idaho [M.S. thesis]: Pocatello, Idaho, USA, Idaho State University, 147 p. Currey, D.R., 1982, Lake Bonneville: selected features of relevance to neotectonic analysis: U.S. Geological Survey Open-File Report 82-1070. Cumming, H.J.K., 1980, Gravity survey of southern Marsh Valley, Bannock County, Idaho [masters’ thesis]: Pocatello, Idaho, USA, Idaho State University, 66 p. Eversaul, M., 2004, Basin Structure of Proposed Late Miocene to Pliocene Supradetachment Basins in Southeastern Idaho Based on Detailed Gravity and Geologic Data [M.S. thesis]: Pocatello, Idaho, USA, Idaho State University, 32 p. Gaylen, R.L., 1978, The Cenozoic Deposits of a Portion of Marsh Valley, Bannock County, Idaho [masters’ thesis]: Pocatello, Idaho, USA, Idaho State University, p. 38. Gilbert, G.K., 1890, Lake Bonneville: U.S. Geological Survey Monograph 1, 438 p. Godsey, H.S., Currey, D.R., and Chan, M.A., 2005, New evidence for an extended occupation of the Provo shoreline and implications for regional climate change, Pleistocene Lake Bonneville, Utah, USA: Quaternary Research, v. 63, p. 212–223, doi:10.1016/j.yqres.2005.01.002. Henrich, W.J., 1979, Gravity survey of northern Marsh Valley, Bannock County, Idaho [masters’ thesis]: Pocatello, Idaho, USA, Idaho State University, p. 54. Janecke, S.U., and Evans, J.C., 1999, Folded and Faulted Salt Lake Formation above the Miocene to Pliocene New Canyon and Clifton detachment faults, Malad and Bannock Ranges, Idaho: Field trip guide to the Deep Creek half graben and environs, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the Geology of Eastern Idaho: Pocatello, Idaho Museum of Natural History, p. 71–96. Janecke, S.U., and Oaks, R.Q., Jr., 2007, Unraveling the Bonneville failure and its aftermath, southeast Idaho: Geological Society of America Abstracts with Programs, v. 39, no. 6, p. 110. Kellogg, K.S., Harlan, S.S., Mehnert, H.H., Snee, L.W., Pierce, K.L., Hackett, W.R., and Rodgers, D.W., 1994, Major 10.2-Ma rhyolitic volcanism in the eastern Snake River Plain, Idaho—isotopic age and stratigraphic setting of the Arbon Valley Tuff Member of the Starlight Formation: U.S. Geological Survey Bulletin 2091, 18 p. Kruger, J.M., Crane, T.J., Pope, A.D., Perkins, M.E., and Link, P.K., 2003, Structural and stratigraphic development of Neogene basins in the Lava Hot Springs and Marsh Valley areas, southeast Idaho: Two phases of extension, in Raynolds, R.G., and Flores, R.M., eds., Cenozoic Systems of the Rocky Mountain Region: Rocky Mountain Section, SEPM, p. 407–458. Link, P.K., and Hodges, M.K.V., 2011, this volume, The Neogene drainage history of south-central Idaho, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, doi:10.1130/2011.0021(05). Link, P.K., and Phoenix, E.C., 1994 (2nd edition 1996), Rocks, Rails and Trails: Pocatello, Idaho, Idaho Museum of Natural History, 194 p. Link, P.K., and Stanford, L.R., 1999, Geologic compilation map of the Pocatello 30 by 60 minute quadrangle, Idaho: Idaho Geological Survey Technical Report 99-02, scale 1:100,000. Link, P.K., Crook, S.R., and Chidsey, T.C., Jr., 1985, Hinterland structure, Paleozoic stratigraphy and duplexes of the Willard thrust system: Bannock, Wellsville and Wasatch Ranges, southeastern Idaho and northern

Neogene drainage development of Marsh and Portneuf valleys Utah, in Kerns, G., and Kerns, R., eds., Orogenic Patterns and Stratigraphy of North-Central Utah and Southeastern Idaho: Utah Geological Association Publication 14, Field Conference Road Log, Day 3, p. 315–328. Link, P.K., Kaufmann, D.S., and Thackray, G.D., 1999, Field Guide to Pleistocene Lakes Thatcher and Bonneville and the Bonneville Flood, southeastern Idaho, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the Geology of Eastern Idaho: Idaho Museum of Natural History, p. 251–266. Long, S.P., and Link, P.K., 2007, Geologic compilation map of the Malad City 30 by 60 minute quadrangle, Idaho, Idaho Geological Survey Technical Report T-07-1, scale 1:100,000. Malde, H.E., 1968, The catastrophic late Pleistocene Bonneville Flood in the Snake River Plain, Idaho: U.S. Geological Survey Professional Paper 596, 52 p. McCoy, W.D., 1987, Quaternary aminostratigraphy of the Bonneville Basin, western United States: Geological Society of America Bulletin, v. 98, p. 99–112, doi:10.1130/0016-7606(1987)98<99:QAOTBB>2.0.CO;2. McQuarrie, N., and Rodgers, D.W., 1998, Subsidence of a volcanic basin by flexure and lower crustal flow: The eastern Snake River Plain, Idaho: Tectonics, v. 17, p. 203–220, doi:10.1029/97TC03762. O’Connor, J.E., 1993, Hydrology, hydraulics, and geomorphology of the Bonneville flood: Geological Society of America Special Paper 274, 83 p. Ore, H.T., 1982, Geologic field guide to Tertiary-Quaternary sediments in the Pocatello region: Northwest Geology, v. 11, p. 47–55. Oriel, S.S., and Platt, L.B., 1980, Geologic map of the Preston 1° × 2° quadrangle, Idaho and Wyoming: U.S. Geological Survey Miscellaneous Investigations Series Map I-1127, scale 1:250,000. Oviatt, C.G., Currey, D.R., and Sack, D., 1992, Radiocarbon chronology of Lake Bonneville, eastern Great Basin, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 99, p. 225–241, doi:10.1016/0031-0182(92)90017-Y. Pierce, K.L., and Morgan, L.A., 1992, The track of the Yellowstone hot spot: Volcanism, faulting, and uplift, in Link, P.K., Kuntz, M.A., Platt, L.B., eds., Regional geology of eastern Idaho and western Wyoming: Geological Society of America Memoir 179, p. 1–53. Pope, A.D., 2002, Geology of the Wakley Peak, Idaho, 7.5′ quadrangle: Multiple phases of late Miocene to Pliocene extension, and relations to the southern Hawkins basin volcanic center [M.S. thesis]: Pocatello, Idaho, USA, Idaho State University, 124 p. Pope, A.D., Blair, J.J., and Link, P.K., 2001, Geologic Map of the Wakley Peak quadrangle, Bannock and Oneida Counties, Idaho: Idaho Geological Survey Technical Report 01-4, scale 1:24,000.

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Reid, T.V., 1997, Gravity-derived subsurface paleomorphology and structure of the lower Portneuf River Valley, Pocatello Idaho [M.S. thesis]: Pocatello, Idaho, USA, Idaho State University, 61 p. Rodgers, D.W., and Othberg, K.L., 1999, Preliminary geologic map of the Pocatello South quadrangle: Idaho Geologic Map 26, scale 1:24,000, 2 sheets. Rodgers, D.W., Hackett, W.R., and Ore, H.T., 1990, Extension of the Owyhee Plateau, eastern Snake River Plain, and Yellowstone Plateau: Geology, v. 18, p. 1138–1141, doi:10.1130/0091-7613(1990)018<1138: EOTYPE>2.3.CO;2. Rodgers, D.W., Ore, H.T., Bobo, R.T., McQuarrie, N., and Zentner, N., 2002, Extension and subsidence of the eastern Snake River Plain, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and magmatic evolution of the Snake River Plain Volcanic province: Idaho Geological Survey Bulletin 30, p. 121–155. Rodgers, D.W., Long, S.P., McQuarrie, N., Burgel, W.D., and Hersley, C.F., 2006, Geologic map of the Inkom quadrangle, Bannock County, Idaho: Idaho Geological Survey Technical Report 06-2, 1:24,000, 2 sheets. Scott, W.E., Pierce, K.L., Bradbury, J.P., and Forester, R.M., 1982, Revised Quaternary stratigraphy and chronology in the American Falls area, southeastern Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 581–595. Shields, R.H., 1978, Depositional Environments of the Sediments in Marsh Valley Near Arimo, Idaho [masters’ thesis]: Pocatello, Idaho, USA, Idaho State University. Steele, E.A., 1980, Depositional Environments of the Sediments in Southern Marsh Valley, Bannock County, Idaho [masters’ thesis]: Pocatello, Idaho, USA, Idaho State University, 58 p. Tauxe, L., Luskin, C., Selkin, P., Gans, P., and Calvert, A., 2004, Paleomagnetic results from the Snake River Plain: Contribution to the time-averaged field global database: Geochemistry Geophysics Geosystems, v. 5, Q08H13, doi:10.1029/2003GC000661. Trimble, D.E., 1976, Geology of the Michaud and Pocatello quadrangles, Bannock and Power Counties, Idaho: U.S. Geological Survey Bulletin 1400, 88 p., scale 1:48,000. MANUSCRIPT ACCEPTED BY THE SOCIETY 9 MARCH 2011

Printed in the USA

The Geological Society of America Field Guide 21 2011

The Neogene drainage history of south-central Idaho Paul K. Link* Department of Geosciences, Mail Stop 8072, Idaho State University, Pocatello, Idaho 83209, USA Mary K.V. Hodges* U.S. Geological Survey, Idaho National Engineering Laboratory Project Office, Idaho Falls, Idaho 83415, USA

ABSTRACT Study of the distribution of the age-populations of detrital zircons in the Snake River system suggest that specific stream systems can be identified based on the detrital-zircon age-population signature (“barcode”) of ancient and Holocene sand deposits. Detrital zircon studies can be used on regional and local scales to determine changes in drainage patterns using both surface and subsurface data. Regional study of drainage patterns using detrital zircons found in Neogene strata of Idaho and southwest Montana suggest northeastward late Miocene to Holocene migration of the Continental Divide from the western side of the Pioneer Mountains to the current position in southwest Montana. Specifically, mixed populations of recycled Proterozoic detrital zircons that define the Wood River drainage are not found in the western Snake River Plain until after 7 Ma. Late Miocene eastward drainage from the central Snake River Plain to southwest Montana is suggested by 9–12 Ma detrital zircons found in fluvial strata less than 6 million years old, of the Sixmile Creek Formation Basalt eruptions of the Eastern Snake River Plain during the Pliocene and Pleistocene also caused drainage diversion. Detrital zircons in Pliocene sands from coreholes at Wendell and Mountain Home Air Force Base contain Big Lost River zircon provenance, suggesting that during the Pliocene, the Big Lost River flowed west along the central Snake River Plain. Late Pliocene and early Pleistocene basaltic volcanoes and rhyolite dome eruptions resulted in volcanic highlands, the Axial Volcanic Zone of the eastern Snake River Plain and the northwest-trending Arco Volcanic Rift Zone (which includes Craters of the Moon volcanic center). The development of these volcanic highlands disrupted the ancestral drainage of the Pliocene Big Lost River system, confining it to the Big Lost Trough, a volcanically dammed basin of internal drainage on the Idaho National Laboratory. After the Big Lost Trough was cut off from the main Snake River, basalt eruptions, local subsidence, and climate controlled the courses of the rivers that flowed into it. Detrital-zircon populations in core samples reveal the provenance of specific sand beds from the Big or Little Lost River systems.

*[email protected]; [email protected]. Link, P.K., and Hodges, M.K.V., 2011, The Neogene drainage history of south-central Idaho, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 103–123, doi: 10.1130/2011.0021(05). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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INTRODUCTION South-central Idaho, north of the Snake River Plain, is an area of superposed late Paleozoic, Mesozoic, Paleogene and Neogene deformation. It also lies on the northern side of the middle Miocene to Holocene Snake River Plain–Yellowstone Hot Spot track. As such, the bedrock geology is varied and the modern topography is scenic and diverse. The two main river systems of this area, the Big Lost River system on the east side of the Pioneer Mountains and the Wood River system on the west side (Fig. 1), presently flow toward the Snake River Plain through Basin and Range valleys. However, the Big Lost River turns northeast upon reaching the Snake River Plain and ends in a playa system (the Lost River Sinks) in a Pleistocene to Holocene volcanically silled basin called the Big Lost Trough (Geslin et al., 1999, 2002). The Wood River on the other hand empties into the Snake River after

cutting narrow canyons through Pleistocene basalt lava flows on the Snake River Plain. In this article and field guide, we summarize evidence from detrital-zircon geochronology that bears on the courses of these streams over the past 10 m.y. In the article by Hodges and Link, we summarize the detrital-zircon signatures or “barcodes” of Holocene streams in the Lost River and Wood River catchments and review detrital-zircon signatures from sediment in drillholes on the Snake River Plain that can be traced to one or the other stream. In the field trip guide, Link traces the path of the Big Lost River from its headwaters in the Pioneer and Boulder Mountains to the Snake River Plain. On the modern drainage divide between the Big Lost and Big Wood rivers there is evidence of Pleistocene drainage capture of the headwaters of Trail Creek, which now flows into the Big Wood River, but which formerly drained into the Big Lost River system.

Figure 1. Map of southern Idaho area showing locations of detrital zircon samples (letters of plots in Figures 2 and 3), locations of coreholes and critical Neoproterozoic provenance areas. Position of hypothesized late Miocene (>7 Ma) Continental Divide is shown as heavy white dashed line and Holocene Continental Divide is shown as solid thin white line. White triangles in the Central and Western Snake River Plain are coreholes sampled for detrital zircons. Dashed circle is Wildhorse Creek in the Pioneer Core Complex, the source for Neoproterozoic detrital grains; Heavy dashed black line is inferred course of paleo–Big Lost River. Black X is Big Southern Butte. Dashed white line is Axial Volcanic Zone and Arco volcanic rift zone. Dashed white line is Big Lost Trough. MHAB is Mountain Home Air Force Base. Location of Beaverhead pluton is shown. Map modified from Link et al. (2005).

The Neogene drainage history of south-central Idaho DETRITAL ZIRCON EVIDENCE FOR NEOGENE DRAINAGE CHANGE IN SOUTH-CENTRAL IDAHO AND PLEISTOCENE DIVERSION OF THE BIG LOST RIVER INTO THE BIG LOST TROUGH, EASTERN IDAHO Mary K.V. Hodges Paul K. Link Introduction Link and colleagues sampled most of the streams in central Idaho (Fig. 1) for detrital zircons and established that zircon agepopulations in 60-grain samples relate to bedrock geology in a predictable manner. Detrital-zircon “barcodes” from small, firstorder drainage systems (Ingersoll et al., 1993) reflect the bedrock geology of fault-bounded Basin and Range systems. In contrast, larger second- and third-order drainage systems, like the Snake River, have a greater number of smaller age-populations representing small point-source magmatic zircons and Proterozoic and Archean zircons recycled through Paleozoic sedimentary rocks into modern streams (Link et al., 2005). Beranek et al. (2006) sampled Miocene strata of the western Snake River Plain and found that the diverse Proterozoic zircon age-populations that characterize the Wood River system only appeared in fluvial deposits <7 Ma. The Pliocene (4–3 Ma) Wood River system drained southward to Glenns Ferry Formation beds at Hagerman Fossil Beds National Monument on the eastern shore of Lake Idaho (near Thousand Springs west of Twin Falls on Fig. 1; Link et al., 2002). Stroup et al. (2008a; 2008b) sampled Oligocene and Miocene fluvial strata of the Renova and Sixmile Creek formations in southwest Montana and found detrital zircon age-populations in rocks >6 Ma that resemble the Wood River signature, but without 45–50 Ma grains from the Challis magmatic pulse. This Sixmile Creek Formation sample also contained a large population of 9–12 Ma volcanic grains likely derived from the central Snake River Plain. The studies by Link, Beranek, and Stroup suggest that until about 7 Ma, the Continental Divide was located in central Idaho southwest of the Wood River drainage, and that the divide has migrated northeastward in the latest Miocene and Pliocene to its present location, concomitant with the migration of Snake River Plain–Yellowstone Hot Spot volcanic centers (Fig. 1). The change in position of the Continental Divide was predicted by studies of the northeast-migrating Snake River Plain–Yellowstone hotspot (Pierce and Morgan, 1992, 2009). Data Central Idaho Detrital-zircon age plots from the Holocene Lost River and Wood River systems (Fig. 1) are shown in Figures 2 and 3 (plots were formerly shown and data are deposited in Link et al., 2005). Detrital zircon U/Pb ages were measured on a sen-

105

sitive high resolution ion microprobe (SHRIMP) at Australian National University (Williams, 1998). Analyses that are more than 10% discordant were eliminated; individual analyses have ~1% precision. Big Lost River System The Big Lost River system, sampled at Lost River Field Station (2A on Fig. 1; LRFS on Fig. 10) and in composite plots at the Highway 26 rest area (2C on Fig. 1; Stop 2-8 on Fig. 10) contains a dominant (75% of the zircon grains) grain population derived from the 45–50 Ma Challis magmatic episode. A smaller (~10%) age-population in the Big Lost River is Neoproterozoic (650–700 Ma). These Neoproterozoic zircon grains have been traced to a 695 Ma pluton in the lower plate of the Pioneer Core Complex, exposed only in the Wildhorse gneiss complex in Wildhorse Creek, a tributary to the Big Lost River (Durk, 2007). In the first-order Big Lost River drainage system, these two populations overwhelm recycled Proterozoic and Archean grains agepopulations and provide a distinctive detrital-zircon barcode for the Big Lost River. Antelope Creek (Figs. 1 and 2B), a tributary to the Big Lost River from the west, drains an area underlain by Challis Volcanic Group, the Mississippian Copper Basin Group of the Copper Basin thrust sheet, and younger strata of the Hawley Creek thrust sheet east of the Copper Basin thrust fault (Skipp et al., 2009). Antelope Creek detrital zircon samples predictably contain abundant Eocene grains and Paleoproterozoic (>1800 Ma) grains recycled through the Copper Basin Group. Since Antelope Creek does not drain the Wildhorse gneiss complex in the lower plate of the Pioneer Core Complex, it lacks the 650–700 Ma grainpopulation that defines the overall Big Lost River provenance. The Little Lost River, another first-order drainage system, (Fig. 2D) is dominated by Challis and Paleoproterozoic detrital zircons. The Paleoproterozoic grains, mainly <1800 Ma, have been recycled through the Mesoproterozoic Lemhi Group exposed in the Lemhi Range (Link et al., 2007; Stewart et al., 2010). Birch Creek (Fig. 2E), which drains the eastern side of the Lemhi Range, contains a unique 500 Ma grain-population (Beaverhead pluton, see Fig. 1) derived from Cambrian and Ordovician alkalic rocks of the Big Creek–Beaverhead magmatic belt (Lund, 2008; Lund et al., 2010). Challis-aged grains are present, but much more sparse than in western streams. A few 40–30 Ma grains (similar to Sample 2I) are present, and represent magmatic activity from southwest Montana. Miocene (<10 Ma) hot-spot– related grains are also present. Figure 2I is a plot from coarse-grained sand and terrace gravel from Antelope Valley south of Challis, near Stop 2-7 of this field trip (see photograph in Fig. 21). The Antelope Valley terrace sample is unique because it has 40–30 Ma zircon grains. The nearest possible source of volcanic rocks of this age is in southwest Montana near Dillon. The gravel containing this zircon population may have been reworked from a post-Challis (45– 50 Ma) and pre–Basin and Range west-draining stream system.

Number

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Link and Hodges

30 Challis Magmatic Event

BIG LOST TROUGH SYSTEM

n=39

0 0

25

50

75

A. Upper Big Lost River 56 grains

100 125 150 Ma

4 2 0

0

500

150

1000

20

1500

2000

2500

3000 Ma

n=37

0 0

25

50

75

B. Antelope Creek 46 grains

100 125 150 Ma

Peace River Arch recycled through Miss. Copper Basin Gp.

0

Number

n=17

150

500

1000

1500

2000

2500

n=9

2

0 3000 Ma

100

n=151

50 0

0

25

50

75

C. Lower Big Lost River 189 grains

100 125 150 Ma Ghost Neoproterozoic

n=38

6 3 0

0

500

150

1000

6 3 0

1500

2000

2500

n=8 0

25

50

75

D. Little Lost River 48 grains

100 125 150 Ma Yavapai-Mazatzal recycled through Belt Supergroup

0

150

500

4

1000

Hotspot volcanics Oligocene volcanics

2

1500

n=40

0

25

50

75

2000

2500

E. Birch Creek 49 grains n=37

Beaverhead Pluton

Figure 4- Page 1 Link et al.

1000

0 3000 Ma

n=12

100 125 150 Ma

500

6 3

0

0 150

3000 Ma

12 6

1500

2000

2500

0 3000 Ma

Figure 2. Detrital-zircon age plots for Lost River system (from Link et al., 2005, fig. 4). Age data are from SHRIMP U-Pb analysis; 60 grain samples. Grains less than 150 Ma are shown in upper left, histogram bin width is 5 m.y. Grains more than 150 Ma are shown in the lower part of each plot; histogram bin width is 25 m.y. Only ages less than 10% discordant are shown. Plots include: (A) Upper Big Lost River at Lost River Field Station; (B) Lower Antelope Creek; (C) Lower Big Lost River, four samples lumped, Stop 2-8 of field trip, including Big Lost River Rest area and Pleistocene deposits directly upstream; (D) Little Lost River, draining area between Lost River Range and Lemhi Range; (E) Birch Creek, draining valley east of Lemhi Range; and (I) Antelope Flat north of Willow Creek Summit (Stop 2-7 of field trip).

Number

The Neogene drainage history of south-central Idaho

14 Challis

7

WOOD RIVER SYSTEM

n=31

0 0

25

50

75

100

125

A. Headwaters of Salmon River 61 grains

150 Ma

150

500

1000

1500

6 3

2000

2

0 3000 Ma

n=11

0 0

25

50

75

100

125

B. Trail Creek 30 grains

150 Ma

Recycled Peace River Arch

Antler-age

0

2500

4

n=30

Number

Yavapai-Mazatzal recycled through Penn-Perm Sun Valley Gp.

Recycled Grenville

0

107

n=19

500

150

1000

1500

2000

2500

2

0 3000 Ma

14 Challis

7

Atlanta Lobe

n=32

0 0

25

50

75

100

125

C. Big Wood River 45 grains

150 Ma

Antler-age

n=13

0

500

150

1000

1500

2000

2500

2

0 3000 Ma

16 Hotspot

8

n=35

0 0

25

50

75

100

125

D. Little Wood River 48 grains

150 Ma

n=13

2 0

0

500

150 50

1500

Atlanta Lobe Idaho batholith

Challis

25

1000

25

50

75

100

125

2500

E. Wood River Composite 184 grains

150 Ma

Recycled Pioneer Core Complex and Sun Valley Gp. Antler-age

0

150

500

Recycled Miss. Copper Basin Gp.

Recycled Grenville

1000

3000 Ma

n=109

0 0

2000

1500

2000

2500

n=75

6 3 0

3000 Ma

Figure 3. Detrital zircon plots (SHRIMP ages) for Wood River system (from Link et al., 2005, fig. 9). (A) Headwaters of the Salmon River. (B) Pleistocene terrace of Trail Creek (below Stop 2-1 of field trip). (C) Lower Big Wood River below Bellevue. (D) Little Wood River below Carey. (E) Lumped Wood River composite signature.

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Wood River System In contrast to the Big Lost River, the Wood River system (Fig. 3; sample locations are 3B, 3C, and 3E on Fig. 1) has a prominent population of Late Cretaceous (95–75 Ma) Atlanta lobe Idaho batholith grains and a broad range of ultimately Laurentian (North American cratonic basement) grains. These include “Grenvillean” age zircons (900–1200 Ma), recycled from the Devonian Milligen Formation and the Pennsylvanian and Permian Wood River Formation of the Sun Valley Group. In the Wood River drainage, Trail Creek and the Wood River (Fig. 3B, 3C, and 3E) contain a 420–440 Ma “Antler age” peak. The Paleoproterozoic (>1800 Ma) age-peak, derived from the Ordovician Kinnikinic Quartzite or the Mississippian Copper Basin Group, is also prominent. The headwaters of the Salmon River (Sample 3A) drain Sun Valley Group strata and predictably contain Challis and mixed Proterozoic <1800 Ma grains. Zircons in the Little Wood River system (Fig. 3D) include Challis-aged (45–50 Ma)) grains and Proterozoic grains recycled from both the Wood River Formation (<1800 Ma) and Copper Basin Group (>1800 Ma). Subsurface Study—Big Lost Trough and Snake River Plain Subsurface SHRIMP analysis of U/Pb ages of detrital zircons in twelve late Miocene to Pleistocene sand samples from six drill cores on the Snake River Plain, Idaho (Hodges et al., 2009), were used to investigate the Pliocene to Holocene courses of the Big and Little Lost River. Locations of wells on the Idaho National Lab are shown on Figure 4. We found that even small 20-grain samples within the Big Lost Trough (Fig. 5A, 5B, 5C, 5D, 6G, 6H) can be attributed to the Big Lost River based on the presence of Neoproterozoic (650–740 Ma, Cryogenian) detrital zircon grains that come from the 695 Ma orthogneiss that intrudes the Wildhorse gneiss in the Pioneer Mountains core complex (Durk, 2007; Durk et al., 2007). These zircon grains traveled down the Wildhorse Creek drainage into the Big Lost River and define the Big Lost River barcode in subsurface Pleistocene sands in the Big Lost Trough. One sample can be traced to the Little Lost River (Fig. 6E). It has a Challis-aged grains and a strong Paleoproterozoic population. It lacks a Neoproterozoic population. There is mixed Big and Little Lost River provenance in ca. 1 Ma (Jaramillo age) sands in corehole C1A (Fig. 6F). Sands from coreholes drilled near Wendell and at Mountain Home Air Base (Fig. 7I, 7J, 7K, 7L) in the central and western Snake River Plain were examined to see if an ancestral Big Lost River signature could be isolated. The Wendell corehole was drilled in the central Snake River Plain as part of the U.S. Geological Survey (USGS) Regional Aquifer System Analysis. In samples from that well, Figure 7I shows data from the sample from depth 137 m, in the Glenns Ferry Formation (Fig. 8) (Whitehead, 1992; Hart and Brueseke, 1999). Figure 7J shows data from 329 m, in beds correlated with the Pliocene or late Miocene middle unit of the Banbury Basalt Formation (Whitehead, 1992). Both of these samples must repre-

sent Pliocene drainage from the Big Lost River, but not the Wood River system, since the both lack Idaho batholith grains. The Mountain Home Air Base Well was drilled to a total depth of 1342 m, but the upper 305 m were not cored (Lewis and Stone, 1988). The sampled strata correlate with the Pliocene Idaho Group, possibly the upper Glenns Ferry Formation (Fig. 8). Figure 7K shows data from the sample taken from 411 m below land surface, which has a zircon assemblage that is very much like that of the integrated drainage of the modern Snake River, with zircons from north and south of the plain, including 155-Ma grains from Jurassic granite in northern Nevada. Figure 7L shows data from the sample taken at 442 m below land surface. It also contains a zircon assemblage that may reflect a northern provenance, with the notable absence of 8–12 Ma central Snake River Plain grains and Jurassic grains. More sampling is obviously needed to gain confidence in these interpretations. Discussion Westward Drainage of Big Lost River In the late Miocene and Pliocene, until about 2.3 Ma, the Big Lost and Big Wood River systems flowed to the west as demonstrated by detrital zircons in cores at Mountain Home. Late Miocene and Pliocene sediment in core at Wendell reflect the Big Lost but not the Wood River system, which must have been farther north. A possible course of the paleo–Big Lost River is shown on Figure 1 (heavy dashed line). Big Lost Trough After about 2.3 Ma the rise of the Axial Volcanic Zone and the Arco–Big Southern Butte Volcanic Rift Zone diverted the Big Lost River system into the Big Lost Trough. Today the Big Lost Trough contains the drainages of the Big Lost River, the Little Lost River, and Birch Creek (Figs. 1 and 4). Detrital zircon studies on Pleistocene sands from coreholes drilled at the Idaho National Laboratory show that eruptions near the Big Lost Trough controlled the deposition of sediments from the streams that flow into it (see cross section, Fig. 9). To the south, corehole C1A yielded three samples (Fig. 6). Sample F (Fig. 6F) is from 402 m below land surface, from a sand bed that lies between normal polarity Jaramillo age basalts and is older than 988 ka. It contains zircons characteristic of both the Big and Little Lost Rivers. Sample G (Fig. 6G) is from 443 m below land surface, and is older than 1.072 Ma, as it is below the normal polarity Jaramillo basalts. Sample G (Fig. 6G) only contains the zircon assemblage characteristic of the Big Lost River. Sample H (Fig. 6H) is from a sand bed 465 m below land surface, is older than 1.37 Ma (Champion et al., 2011), and also has the Big Lost River zircon assemblage without admixture from other streams. Before 1.072 Ma, the Big Lost River flowed into the C1A area, and sometime between 1.072 Ma and 988 ka, the obstruction that prevented the Little Lost River from contributing sediment to that basin was overtopped, and sediment from the Big and Little Lost River was mixed (Hodges et al., 2009; Champion et al., 2011).

The Neogene drainage history of south-central Idaho

109

113°

112°30'

EXPLANATION

44°

ch Bir

Approximate area of volcanic rift zones (VRZ) and axial volcanic zone Corehole from which samples taken for paleomagnetic inclination analysis Corehole from which samples taken for detrital zircon studies Location of field trip stop Big Lost Trough

Mud Lake

tte Bu

Mud Lake

-K

Birch Creek sinks TAN CH1

te VR

’s lf A

RA

COREHOLE 2-2A

Ha

R

VE

Z

ell Birch Creek sinks

R

RI

i v er

33

ut

GIN 5

eH

ost

ST

L

Terreton

B le

ett

GIN 6 TAN

TAN CH2

idg

LO

Li ttl e

Howe

r ula

Radioactive Waste Management Complex Advanced Test Reactor Complex Idaho Nuclear Technology and Engineering Center Central Facilities Area Naval Reactor Facility Test Area North Materials and Fuels Complex

R va La

RWMC ATRC INTEC CFA NRF TAN MFC

126A

ek Cre

Facility with location identifier

rc Ci

E I NG MH RA LE

Cross section line

cre

E

NG

Big Lost River sinks and playas

Z

VR

we

Ho NRF 7P

-E

Arco 134

B ig

Lost

43° 30'

133

ATRC

121

C’

INTEC

ICPP 023 123

Z

Middle 1823 Middle 2050A ICPP-214

VR

A-BSB VRZ

te

26 20

ut

tB

as

NRF

ANL-OBS-A-001 MFC

NPR TEST/W-02

ANL DH 50

20

128

Middle 2051

R iver

CFA

C1A

129

127 130 131

STF-AQ-01

East Butte

ARA-COR-005

Middle Butte

RWMC

C 135

IDAHO

132 IDAHO NATIONAL LABORATORY BOUNDARY

Eastern Snake River Plain

COREHOLE 1

SPREADING AREAS

Big Southern Butte

e

on

Z nic

a

olc

Atomic City

lV xia

A

26

Idaho National Laboratory Boise

Idaho Falls Pocatello

Twin Falls

0 0

5

10 5

15 KILOMETERS 10

15 MILES

Base from U.S. Geological Survey digital data, 1:24,000 and 1:100,000 Universal Transverse Mercator projection, Zone 12 Datum is North American Datum of 1927

Figure 4. Map of Idaho National Laboratory and vicinity. Dots with red hexagons are drill holes (USGS 134, Middle 1823, Middle 2050A, and C1A) sampled for detrital zircons (Figs. 5, 6, 7) and shown in cross section in Figure 9. Idaho National Laboratory boundary shown by gray dashed and dotted line. Other features include the Big Lost Trough (heavy dashed black line), and shaded areas for the Axial Volcanic Zone and Arco–Big Southern Butte Volcanic Rift Zone (A-BSB VRZ). Map modified from Hodges et al. (2009).

110

Link and Hodges

All sands A through D have Big Lost River provenance Challis magmatic event

8 4

59PL05--USGS 134, 940-949 ft. 20 grains

n=14

A

0 0

50

100

150 Ma

Neoproterozoic (ghost) grains 0

500

1000

1500

8

2000

4

n=6

2 0 3000 Ma

2500

B

60PL05--Middle 2050A--742-743 ft. 23 grains

n=15

4 0 0

50

100

150 Ma

2

n=8

1

0

500

1000

8

0 3000 Ma

1500

2000

2500

n=13

61PL05--Middle 2050A, 1056-1061 ft. 20 grains

C

4 0 0

25

50

75

100

150 Ma

125

4

n=7 2

0

500

1000

1500

16

2000

0 3000 Ma

2500

D

62PL05--Middle 1823--1311-1312 ft 35 grains

n=24

8 0 0

50

100

150 Ma

n=11 2

0

500

1000

1500

2000

2500

0 3000 Ma

Figure 5. Probability-frequency plots for SHRIMP U-Pb ages for detrital zircon grains in Neogene to Holocene Idaho National Laboratory drillhole sands. (A) USGS 134 287–289 m; (B) Middle 2050A, 226 m; (C) Middle 2050A, 322– 323 m; (D) Middle 1823, 400 m. Grains younger than 150 Ma are plotted in the upper part of diagram, and grains older than 150 Ma are shown in the lower half. Definitive populations are labeled. Locations of wells are shown in Figure 4. Figure modified from Hodges et al. (2009).

The Neogene drainage history of south-central Idaho

Challis magmatism

111

E Little Lost River signature Strong Paleoproterozoic peak

F Mixed, provenance, with significant input from Little Lost river Paleoproterozoic grains

G Big Lost River provenance

H

Figure 6. Probability-frequency plots for SHRIMP U-Pb ages for detrital zircon grains in Neogene to Holocene Idaho National Laboratory corehole sands. Labeling of plots is sequential from Figure 5. (E) Middle 1823, 450 m; (F) C1A, 402 m; (G) C1A, 443 m; (H) C1A, 465 m. Locations of coreholes are shown in Figure 4. Figure modified from Hodges et al. (2009).

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Twin Falls Strong population Challis grains volc. fieldhotspot 7-15 Ma Idaho batholith Challis

67PL05 Wendell core--137 m (450 ft.)--101 grains

I

I

Northern side of SRP Idaho batholith

provenance 6 grains 545 Recycled Grenville, Mesoproterozoic 6 Neoproterozoic to 670 Ma and Paleoproterozoic grains Recycled Proterozoic and Archean grains 545-675 Ma

3 grains 7-11 Ma

Strong Challis Challis peak

J

J

Northern side of SRP Strong Grenville, scatteredprovenance

One 3 grains 168 Ma 5703tograins 670 Ma grain 570 -670 Ma

Y-SRP grains

68PL05 Wendell core--329 m (1090 ft.)--60 grains

Strong Grenville and Paleoproterozoic Mesoproterozoic, scattered Mesoproterozoic and Paleoproterozoic

Challis

Idaho batholith

51PL06 MHAB core--411 m (1349 ft.)--60 grains

K

K

c Integrated Snake River barcode 6 grains 152-163 Ma Jurassic from Jurassic from northern northern Nevada Nevada 3 grains 614 +/- 7 Ma

Integrated provenance from north and Scattered Proterozoic south sides of SRP

Scattered Proterozoic

150

4 grains 3.5 to 4.7 Ma 6 grains 14.4-16.5 Ma

No Jurassic!

Scattered Proterozoic

52PL06 MHAB core--442 m (1450 ft.)--25 grains

LL

Northern provenance? Pliocene grains from Magic Reservoir volcanic field

Figure 7. Probability-frequency plots for SHRIMP U-Pb ages for detrital zircon grains in Neogene corehole sands from Wendell and Mountain Home coreholes. Labeling of plots is sequential from Figure 6. Locations of coreholes shown in Figure 1. (I) Wendell core, 137 m; (J) Wendell core 329 m; (K) MHAB (Mountain Home Air Base) core 411 m; (L) MHAB core 442 m. Figure modified from Hodges et al. (2009). SRP—Snake River Plain.

The Neogene drainage history of south-central Idaho

Eastern SRP (INL wells) Central SRP Western SRP (Wendell well) (Mountain Home USGS 134 Air Base well) Middle 1823 Middle 2050A Polarity Chron Polarity (black=normal, C1A

Rock Units

D E

2

Paleo-Big Lost River disrupted by

4

5

Pliocene

J (older than I, younger than 7.1 Ma)

K, (younger than L)

L, (younger than 4.2 Ma)

West flowing paleoBig Lost River drains north side of eastern SRP and runs through Wendell and Mountain Home Air Base into Lake Idaho

Banbury Basalt Fm2

(younger than J, may be younger than 3.5 Ma)

Lake Idaho (WSRP)

Paleomagnetic data not available for Wendell and Mountain Home Air Base cores

Idaho Group

Axial Volcanic Zone

I,

Yellowstone Plateau4

Big Lost TroughNorthern streams cut off from main Snake River by Arco Rift and Axial Volcanic Zone (AVZ)

1.77-1.95 Ma (Olduvai)

3

Hotspot Volcanic Fields

Magic Reservoir 2 Heise1

Matuyama

A C F G H

Snake River Group (basalt flows and thin fine-grained sediments)

0.99-1.07 Ma (Jaramillo)

B

Glenns Ferry Fm3

0.78 Ma (Brunhes -Matuyama)

Pleistocene

Landscape development

white=reversed)

Brunhes

1

Middle 2050A yielded two samples, Figure 5B and 5C. Sample 5B is from 226 m below land surface, is found in sand between normal polarity basalts that are slightly younger than 780 ka (Fig. 9). Sample 5C is from a sand bed that lies between normal polarity basalts that are older than 988 ka and younger than 1.072 Ma. Both Sample 5B and 5C have Big Lost River zircon assemblages; the Little Lost River was excluded from the Middle 2050A area at the times of their deposition (Hodges et al., 2009; Champion et al., 2011). Sample A (Fig. 5A) came from the bottom of corehole USGS 134 and contains the Big Lost River zircon assemblage (Hodges

Tuana Gravel Fm5

Epoch

Age, Ma

In Middle 1823, two samples were recovered from below 1.37 Ma basalts. The younger sample, Figure 5D, is from 400 m below land surface, and contains sand from the Big Lost River. Sample 6E does not have enough zircon grains to make a statistically robust interpretation, but it is probably from the Little Lost River. If sample 6E is from the Little Lost River, before 1.37 Ma, the Middle 1823 basin was accepting sediment from the Little Lost River, with progradation of Big Lost River sediment after that. Around 1.37 Ma, the river valley filled with basalt (Hodges et al., 2009; Champion et al., 2011).

113

Figure 8. Phases of drainage development on the Snake River Plain (SRP). Figure shows stratigraphic names, age constraints, magnetic polarity, and stratigraphic locations of detrital-zircon samples. Transition from through-flowing paleo–Lost River to initiation of the Big Lost Trough is indicated by thick dashed line. General age constraints are after Pierce and Morgan (1992). Magnetic polarity data from Champion et al. (2011). Other references: 1—Morgan (1992) and Morgan and McIntosh (2005); 2—Bonnichsen and Godchaux (2002); 3—Hart and Brueseke (1999); 4—Christiansen (2001); 5—Sadler and Link (1996); 6—Tauxe et al. (2004, table 4 and fig. 12). Figure is modified from Hodges et al. (2009).

114

Link and Hodges

Southwest

55° -70° -62° -64°

NPR Test /W-02

Middle 2051

Middle 2050A

1,400

C1A

1,500

A′ Middle 1823

km sca 7.3 ot to (n

USGS 134

le)

Land surface

VERTICAL EXAGGERATION 30X

ALTITUDE, METERS ABOVE THE NGVD 29

Northeast

A

) km cale 4.3 t to s (no

1,600

54° 68° 60° -6 6 ° -6 9 °

-6 1 °

1,300

-7 1 °

A

-7 2 ° -7 0 ° -7 1 ° -7 1 °

1,200

52° 54° 53° 55°

Post Jaramillo

54° 58° 52°

1,100

-7 0 °

F

Jaramillo

- 63° 51°

5 2° 7 2° -6 7° -6 8° - 6 9° 5 3° 5 2° 5 1° 58°

5 7°

44° 45°

459° 8° 5 3° -6 1° -5 7°

4 9°

G Matuyama 1.37 Ma H

-6 3° -5 7°

-6 6° -6 9°

-6 7 °

D E

B Early Basa

-68°

Matuyama ?

-38°

?

-54°

C

-67°

Matuyama 1.21 Ma

-23° 5° -69°

?

Unrecovered Sediment and basalt -55°

-63°

-6 6 °

-70°

-4 9 °

Pre-Olduvai

900 *Note: Corehole NPR Test/W-02 was drilled to 1,524 m, but only sampled to 1173 m

5

10

Olduvai

70°

Olduvai Lakebeds

0 KILOMETERS

Uncorrelated

-65°

-6 3 °

800

Matuyama 1.256 Ma

-61° 1 1°

-4 2 ° -6 7 °

-6 6 °

Uncorrelated

-59°

?

-66

55°

Uncorrelated

Post Olduvai 1,000

l Brunhes

5 4° -7 2° -6 9° -6 9° -6 8° -6 8° -6 9° -6 8° 5 5° 5 5° 5 3°

-7 0 °

-61° -64° -64° -66°

-55° -56° -57° -46°

15

20

700 0 MILES

5

10

EXPLANATION Brunhes Normal Polarity Chron-Matuyama Reversed Polarity Chron boundary (0.78 Ma to Holocene)

Olduvai Normal Polarity Subchron boundary (1.945 to 1.778 Ma)-

Jaramillo Normal Polarity Subchron Boundary (1.072 to 0.988 Ma)

Water table, approximate

Sediment USGS 134 Sample A 287-289 M BLS Middle 1823 Samples D 400 M BLS E 450 M BLS

C1A Samples F 402 M BLS G 443 M BLS H 465 M BLS Middle 2050A Samples B 226 M BLS C 322 M BLS

Figure 9. Cross section A–A′ modified from Champion et al. (2011) showing subsurface basalt stratigraphy (older than 0.780 Ma, cut away between Middle 2051 and Middle 1823 to show USGS 134) beneath the Big Lost River Rest area, Stop 2-8.

The Neogene drainage history of south-central Idaho et al., 2009). USGS 134 is the most northerly of the coreholes in this study, and the sample is from a sand bed below Jaramillo age basalts that are older than 988 ka. USGS 134 was not drilled through the sediment layer where the sample was taken so the age of the sediments cannot be further constrained (Champion et al., 2011). Conclusions The age populations of detrital zircons found in first-order streams in central Idaho are a reliable indicator of the bedrock geology of the drainage basins of those streams. The presence or absence of uniquely aged zircons in downstream sand deposits can identify the streams from which they came, and, if independent geochronology is available, can be used to determine the timing of drainage disruptions. Detrital zircon studies in central Idaho show that the interaction of active extensional tectonics, surface uplift and subsidence due to isostatic loading of the Snake River Plain, plus construction of volcanic domes and lava fields, all have interacted to modify the courses of Neogene to Holocene drainage. In the middle Miocene, the headwaters of the Big Lost and Wood rivers in the Pioneer and Boulder Mountains presumably drained northward or eastward into the Missouri River system (Stroup et al., 2008a). Recycled Paleoproterozoic detrital zircons, sourced east of the Idaho batholith in the Wood River drainage are not found in western Snake River Plain sediments until after 7 Ma (Beranek et al., 2006). However, to date, late Miocene fluvial strata containing Big Lost River and Wood River detrital zircon barcodes to the north in Montana have not been found. Indeed, the Sixmile Creek Formation near Dillon, Montana, contains Paleozoic zircons most likely from the Big Wood River system, but lacks Challis volcanic grains which should form the largest age-population (Stroup et al., 2008b). The Pliocene Wood River drained southwestward into the eastern side of Lake Idaho and the Snake River system. Since sediments with Challis-aged zircons that would demonstrate the hypothesized southern outlet of a late Miocene Wood River into northern Nevada have not been found (i.e., Link et al., 2002; 2005), our preferred hypothesis is that the >7 Ma Wood River flowed northeast. Sample 7K demonstrates that in the Pliocene, northern Nevada drainage was northward from 155 Ma Jurassic plutons to Mountain Home during deposition of the Glenns Ferry Formation Today the Wood River flows southwestward into the Snake River, whose ultimate base level is the Pacific Ocean. On the field trip at Stops 2-1 and 2-2, we will observe Pleistocene stream capture by Trail Creek, tributary to the Wood River, of what was formerly the head of the Big Lost River. The Pliocene Big Lost River drained westward through Wendell and Mountain Home. In the Pleistocene, the Big Lost River was diverted so it now flows northeastward into the Big Lost Trough, with a base level controlled by the playa surface there.

115

FIELD TRIP GUIDE: THE COURSE OF THE BIG LOST RIVER Paul K. Link Geology of East-Central Idaho The northwest-striking Wood River, Big and Little Lost River and Birch Creek valleys (Fig. 10) are parallel to the structural grain of the Basin and Range province. The upper part of the Big Lost River, on the other hand, consists of several northeast-flowing forks that cross northwest-striking structures that separate different bedrock geology. The area drained by the North Fork of the Big Lost River and Summit Creek contains Mesozoic thrust faults and Paleogene high- and low-angle normal faults (Rodgers et al., 1995; Link et al., 1995, 1996; Skipp et al., 2009). From west to east, the Boulder and Pioneer Mountains contain the Pioneer thrust plate, intruded by the Atlanta lobe of the Idaho batholith on the west; the Copper Basin thrust plate that contains the thick, dark-colored Mississippian Copper Basin Group above Lower Paleozoic platform strata including the Ordovician Kinnikinic Quartzite; and the Hawley Creek thrust plate that contains Paleozoic continental platform strata of the west-facing Laurentian continental margin (Link and Janecke, 1999; Skipp et al., 2009). Wildhorse Creek, tributary to the East Fork of the Big Lost River, drains Archean orthogneiss and Mesoproterozoic metamorphic rocks of the Wildhorse gneiss complex of the Pioneer Mountains core. The Wildhorse gneiss is intruded by a unique Neoproterozoic (695 Ma) alkalic pluton (Durk, 2007) and by ca. 48 Ma Eocene granitic stocks of the Pioneer Intrusive Suite. The Pioneer Intrusive Suite is sheared along the top-to-thewest Hyndman shear zone, a ductile normal fault. The Hyndman shear zone is visible on the north face of the Devil’s Bedstead (Fig. 11), and separates Neoproterozoic and Lower Paleozoic quartzite and marble of the upper part of the metamorphic core from the Archean to Neoproterozoic Wildhorse gneiss complex, all structurally below the Wildhorse detachment fault (Dover, 1981, 1983; Link et al., 2010). The Wildhorse detachment fault (Fig. 10; seen from Stop 2-3; Fig. 11) separates Paleozoic sedimentary rocks of the Pioneer thrust sheet above from Paleozoic quartzite and marble (part of the Copper Basin thrust sheet) and Wildhorse gneiss complex below. Movement direction across the Wildhorse detachment on the northwest side of the core complex was uniformly top to the west-northwest (295 ± 15°; Diedesch and Rodgers, 2010). Uplift of the Pioneer Mountains is mainly post-Oligocene (Vogl et al., 2010). The Summit Creek stock, in the upper plate of the core complex along Summit Creek east of Stop 2-1 of this field trip, was proposed to be the beheaded top of the Pioneer Intrusive Suite (Wust, 1986). Rodgers et al. (2002) detail the formation of Neogene eastern Idaho sedimentary basins as controlled both by Basin and Range extensional tectonics and Snake River Plain uplift and

115°

2-1

2-2

Lit

r ve

B ig d

oo W Ri

Low-angle normal fault

Wildhorse Detachment

. l Cr Trai Summit Ck.

Boulder Mtns.

N. Fk. Big Lost River.

2-4 2-3

2-6 LRFS

2-5

dR

t r ive

114°

Carey

oo le W

An

pe lo te

. Cr

113°

2-8

Lost River Sinks

Howe Pt

MT ID

Figure 10. Location map showing main geologic structures and field trip stops (numbered 2-1 to 2-8). Fault segment boundaries on range-bounding normal faults are indicated by heavy arrows and are named as follows. Lost River fault: A—Arco; C—Challis; MK—Mackay; PC—Pass Creek; TS—Thousand Springs; WS—Warm Springs. Lemhi Fault: E—Ellis; FS—Fallert Springs; H—Howe; G—Goldburg; M—May; S—Sawmill Gulch; WC—Warm Creek. Beaverhead fault: BD—Blue Dome; BM—Baldy Mountain; L—Leadore; MG—Mollie Gulch; N—Nicholia. Map modified from Link and Janecke (1999, fig. 2). LRFS—Lost River Field Station; MT—Montana; ID—Idaho.

44°

2-7

Wildhorse Ck.

116 Link and Hodges

The Neogene drainage history of south-central Idaho subsidence. Southward tilting of the surface adjacent to the northern Snake River Plain occurred in the Carey area before 8 Ma (Michalek, 2009). Subsidence of the Plain and uplift of adjacent ranges may be accommodated by lower crustal flow away from the subsiding Plain (McQuarrie and Rodgers, 1998), which itself is a response to loading by a mid-crustal gabbroic sill beneath the Plain (Peng and Humphreys, 1998). Neogene Drainage Change Eastern Snake River Plain drainages have a complex pattern, and are not “well-adjusted” to topography, a fact pointed out by Hayden (1883), emphasized by Ore (1999), and documented using age-populations of detrital zircons by Link et al. (2005), Beranek et al. (2006), Stroup et al. (2008a; 2008b) and Hodges et al. (2009). In addition to being dammed by lava flows, Neogene drainages and lake systems of the Intermountain West have been controlled by development of the Basin and Range province and the northeastward migration of the Yellowstone–Snake River Plain Hot Spot, with its hypothesized topographic bulge (Pierce and Morgan, 1992; 2009). The main theme of this field trip is that in the Pleistocene, the head of the Big Lost River was beheaded by Trail Creek, tributary to the Big Wood River. The ultimate cause of this drainage capture is that the Pleistocene base level of the Big Lost River is the Lost River Sinks (Fig. 10), whereas the Wood River flows, via Malad Gorge east of Bliss, into the Snake River and ultimately into the Pacific ocean. The timing on this capture is not constrained, though we suggest an early Pleistocene age because glaciated Park Canyon (Stop 2-2) seems to require a glacier fed by what is now the headwaters of Trail Creek on the northeast side of the Boulder Mountains. Field Trip Stops From Sun Valley proceed northeast up Trail Creek Road to the summit. From Mackay drive north on Highway 93 to the Summit Creek/Trail Creek road to Sun Valley, turn west and drive to the summit parking lot. Stops are numbered 2-1, 2-2, etc., because this is the second day of a larger field trip for this meeting (Thackray et al., this volume). Stop 2-1. Crest of the Boulder and Pioneer Mountains (11T 720550E, 4855740N, 2537 m, 7678 ft) Stop 2-1 is on the crest of the Boulder and Pioneer Mountains, on the divide between the west-flowing Big Wood River and east-flowing Big Lost River. To the west, near Sun Valley and Boise, is the interior of the central Idaho Sevier orogenic belt, where the Late Cretaceous Idaho batholith intrudes the Wood River–Milligen structural stack within the Pioneer thrust plate (Link et al., 1995). To the east are thrust faults cutting lower marine Paleozoic strata. The Pioneer thrust fault is located in Little Fall Creek just northwest of Stop 2-2.

117

Here, the even northwestward gradient of Summit Creek and Park Creek, tributaries to the Big Lost River, are cut by the steep southwestward gradient of Trail Creek, tributary to the Big Wood River. This abrupt escarpment is vividly exposed in the canyon to the right (Fig. 12). At mile 0.0, drive northeast from Stop 2-1. At mile 1.3 at Park Creek sign, turn left (northwest) just before campground. Continue to mile 2.2 and stop in meadow just before gentle summit. Stop 2-2. Head of Park Creek (11T 719440E, 4857320N, 2340 m, 7734 ft) The bedrock here is imbricated Devonian to Ordovician finegrained mudrock of the Milligen, Trail Creek, and Phi Kappa formations. On the Park Creek side of the divide, the gradient is low in the formerly east-flowing glaciated valley (Fig. 13). To the west over the gentle summit, Trail Creek drops southward via waterfalls into the main Trail Creek canyon. Turn around and proceed back to Park Creek campground. At T intersection (mile 6.3) turn left on Summit Creek Road. Proceed northeastward and structurally lower, crossing Pioneer thrust fault and into the underlying Copper Basin thrust plate near the Little Fall Creek turnoff. Continue to mile 11.1. Stop 2-3. View of Confluence of North Fork, East Fork, Summit Creek and Kane Creek to Make the Big Lost River (11T 726890E, 4863330N, 2183 m; 7193 ft) Here is the confluence of several forks of Big Lost River (Fig. 11). Glacial geology in this area has been studied by Evenson and colleagues over three decades (Evenson et al., 1982, 1987; Barton, 2007). The Kane Creek, Summit Creek, and the North Fork of the Big Lost River glaciers coalesced into the North Fork glacier complex. Glaciers from Wildhorse Creek formed a lower complex. Pinedale glaciers (Wildhorse Canyon and Devils Bedstead advances, 35–17 ka) in this confluence repeatedly clogged drainages, producing Glacial Lake East Fork in the East Fork of the Big Lost River. This ice-dammed lake flooded at least once, at about 17.3 ka (Barton, 2007), with boulders (i.e., Fig. 14) rafted by ice as far as Box Canyon on the Snake River Plain just upstream from Stop 2-8 (Rathburn, 1993). Bedrock geology here is Mississippian Copper Basin Group unconformably overlain by Eocene Challis Volcanic Group. A prominent north-striking normal fault cuts the front of the mountain to the southeast, behind the East Fork of the Big Lost River. Continue east on Trail Creek/Summit Creek road, passing at mile 12.3 Kane Creek Road on right. At mile 12.8, cross bridge over North Fork Big Lost River. At mile 15.2, pass major intersection with Copper Basin Road to right. At mile 20.1, pass Bartlett Point Road to right. At mile 24.8, pass Walker Way to left. The Swenson Butte ice-rafted boulder is in the late Pleistocene outwash terrace to the left (Fig. 14; Barton, 2007). At mile 28.1, turn right on unimproved road to West Chilly Butte, and stop at the base of the slope in 0.2 miles.

118 E

Link and Hodges Wildhorse Detachment

Hyndman Shear Zone

Devil’s Bedstead

W

SE

NW

Bald Mtn.

Figure 11. View south up Kane Creek to Devils Bedstead peak. Bedstead is underlain by Proterozoic and Paleozoic paragneiss which are part of the lower plate of the Pioneer Core Complex. The Wildhorse Detachment Fault is located in prominent saddle on left (east) side of peak. Mississippian Copper Basin Group lies above the fault. Late Pleistocene glacial moraines and inset outwash terraces fill the valley.

Stop 2-4. View of Borah Peak Looking Northeast from West Chilly Butte (12T 265640E; 4881990N; 1935 m, 6526 ft) The bedrock here is folded Mississippian White Knob Limestone. We are in the east-tilted hanging wall of the Thousand Springs segment of the Lost River fault (Fig. 10), which here has 2.7 km of structural relief and a slip rate of <0.1 m/ka over the past 100,000 years (Scott et al., 1985). The panoramic view to the east extends from the Pahsimeroi Mountains on the north, southward over Willow Creek Summit, to Borah Peak, and then south to Leatherman Peak, Mount McCaleb and Mackay (Figs. 15, 16). The bedrock of the Lost River Range consists of folded

NE

SW Eocene Summit Creek Stock

Figure 12. View southwest down the canyon of Trail Creek. Bald Mountain, the main Sun Valley ski hill, is labeled in the distance. This canyon is steep and glaciated in the upper part. The headwaters of Trail Creek enter the canyon from the west (right) over a several hundred foot waterfall. It is this creek (the head of Trail Creek) that we assert formerly flowed eastward into what is now the Big Lost River drainage.

but mainly east-dipping Ordovician through Permian carbonate and quartzite strata that are part of the Hawley Creek thrust plate. The lower Mississippian McGowan Creek Formation forms a prominent talus slope below cliffs of the upper Mississippian carbonate back complex (Skipp et al., 2009). K.L. Pierce and Scott (1982) pointed out that the alluvial fans on the west front of the Lost River Range have been largely inactive in the Holocene, and were fed by late Pleistocene streams with seasonal flows at least ten times larger than present discharge. Patterson (2006) studied Holocene and Pleistocene deposition on alluvial fans of the active Mackay segment of the Lost River fault, in the distant right part of this view. Ages

SW

Figure 13. View southeast down the head of Park Creek, tributary to Summit Creek and the Big Lost River. This headwater drainage is dry, and has no catchment. The catchment has been captured by the westflowing Trail Creek that directly behind this view drains into a waterfall and the Trail Creek Canyon.

Folded Miss. White Knob Ls.

NE

Figure 14. The Swenson Butte boulder of Wildhorse gneiss rafted into Big Lost River valley about 17.3 ka by a late-glacial flood from Glacial Lake East Fork (cosmogenic isotope date by Barton, 2007).

The Neogene drainage history of south-central Idaho NW Borah Peak

SE

119 Borah Pk. Horst Fault

NW Doublespring Pass

Rock SE Creek

Leatherman Peak Fault Scarp North Chilly Butte

Lost River Field Station

Figure 16. View to northeast of Doublespring Pass and central and northern Chilly Butte. Unnamed mountain in central part of view is underlain by Mississippian strata. Borah Peak fault scarp is shown in foreground. The northern Borah Peak horst normal fault is indicated on right in Rock Creek. North Chilly Butte in foreground.

Figure 15. Big view to southeast of Lost River Range from west Chilly Butte. Borah Peak fault scarp is clear along the base of the mountains. High peak is Borah Peak.

of surfaces were calculated at <43 ka based on thickness of carbonate coats. Small steep Holocene debris-flow fans are deposited on top of incised inactive Pleistocene sheetflood-dominant alluvial fans. J.L. Pierce, students, and colleagues have been investigating the ages of the fan surfaces on the hanging wall of the Lost River fault (Sutfin et al., 2010; Kenworthy et al., 2010). About 88% of fan area was deposited during the past 70 ka, with less than 10% of the fan area active in the last 10,000 years. The surfaces of the Ramshorn fan, east of Moore in the southern Big Lost River Valley, yield concordant 230Th/U and optically stimulated luminescence (OSL) ages of 40–15 ka. The fans were thus active during both late Pleistocene glacial (oxygen isotope stage [OIS] 2 and 4) and cool interglacial (OIS 3) times, and have been largely inactive during the warmer Holocene interglacial.

SE

North Chilly Butte

Turn around and proceed back to Trail Creek Road; turn right. At mile 30.4, cross intersection with Old Chilly Road. At mile 32.5, turn left at T intersection with Highway 93. Proceed north toward Challis. At mile 36.9, pass Borah Peak trailhead junction to right. At mile 37.1, pass Dickey Golf Course Clubhouse on left (Fig. 17). Stop at mile 38.5 at the Borah Peak Historical Marker. Stop 2-5. View of Borah Peak Fault Scarp and Doublespring Pass (12T 267620E; 4891680N; 1945 m; 6372 ft) The Borah Peak fault scarp is obvious low in the view to the east. On 28 October 1983, a magnitude 7.3 earthquake caused a line of surface faulting 34 km long and vertical offset of 2 m just east of here. Several guidebook articles describe the 1983

N

Borah Peak

S

Chicken-out Ridge

NW Fault Scarp

Figure 17. Dickey Golf Course Clubhouse. Chilly Slough and northern Chilly Butte in background.

Figure 18. Cedar Creek and Borah Peak horst. Borah Peak fault scarp is plain at the base of the mountains. Cedar Creek contains multiple generations of late Pleistocene moraines, cut by the Borah Peak fault.

120 NE

Link and Hodges Borah Peak Axis of syncline

SW

S

N Terrace

Figure 19. View to southeast of Borah Peak horst.

Borah Peak earthquake and fault features (Crone, 1987; Link and Janecke, 1999). A short distance to the north up the Doublespring Pass Road is a geological site where the fault is exposed cutting Pleistocene alluvial fans. The talus-covered mountain to the north is underlain by mainly Mississippian strata (Fig. 16). The McGowan Creek Formation forms the talus and the carbonate bank assemblage makes up the cliffs at the top of the hill (Ross, 1947; Janecke and Wilson, 1992). The steep canyons of Cedar Creek drain the west side of Borah Peak. A north-striking Eocene normal fault through Rock Creek forms the north side of the Borah Peak horst (Fig. 16). Lower Paleozoic strata are folded into a prominent syncline along the main ascent route for Borah Peak, known as Chicken-Out Ridge (Fig. 18). The brightly colored rock is Devonian Grandview Dolomite (Grader and Dehler, 1999). Continue north on Highway 93 to mile 45.6 at Willow Creek summit. Stop 2-6. Willow Creek Summit (12T 261940E; 4902540N; 2080 m; 7086 ft) This is the divide between the Big Lost River and the Salmon River to the north. The bedrock is folded and poorly exposed Mississippian White Knob Limestone. Figure 19 shows the view to Borah Peak and the Lost River Valley to the south.

W

Figure 20. Terrace of Antelope Flat south of Grandview Canyon. Detrital zircon sample shown in Figure 2I comes from this terrace.

Sheedy (2003) mapped the glacial landforms and normal faults in the Willow Creek Summit area. She also analyzed detrital zircons from several Pleistocene and older fluvial deposits in Round Valley to the north. Although our model (Hodges and Link, Fig. 1) predicts that the Big Lost River drained northward into the Salmon River in the late Miocene, no high-level gravels containing distinctive Pioneer core complex zircons have been found. There are very few highlevel erosional benches or depositional terraces. The highest and possibly oldest terrace (Fig. 19) was sampled just south of Grandview Canyon. It contains an unusual detrital zircon assemblage (Fig. 20) more characteristic of southwestern Montana than east-central Idaho. Proceed downhill to the north toward Challis. At mile 48.3, pass Road Creek gravel road on left. At mile 50.9, pass Spar Canyon road on left. Enter area known as Antelope Flat. At mile 52.1 on the left is a view of a terrace with Oligocene detrital zircons (Fig. 20). Detrital zircons from this terrace are shown in Figure 2I. View north of terrace and dipping Challis Volcanic Group draped over Devonian carbonate strata. Continue on highway through Grand View Canyon and stop at mile 57.5 at the north end of Grand View Canyon.

E N

S Miss. Surrett Cyn Ls.

Figure 21. North end of Grandview Canyon, superposed onto Devonian dolomite.

Figure 22. Number Hill in Arco, Idaho. High School graduating classes have painted their graduation year here for 90 years. Bedrock is Mississippian carbonate of the Scott Peak, South Creek, and Surrett Canyon formations.

The Neogene drainage history of south-central Idaho Stop 2-7 Superposed Grand View Canyon in Jefferson Formation (11T 732910E; 4916830N; 1793 m; 6498 ft) We have driven northward and topographically downward through a superposed canyon cut in black dolomite of the Grandview Member of the Devonian Jefferson Dolomite (Fig. 21). Such a canyon demonstrates Neogene base level fall, tied to the Salmon River canyon to the north. On the left across the highway is a dirt track that provides access to a bioherm/biostrome complex in the dark dolomite member of the Jefferson Dolomite (Isaacson and Dorobek, 1988; Isaacson et al., 1988; Grader and Dehler, 1999). Turn around, and head back through Mackay and Arco. Coming into Arco folded Mississippian carbonate bank strata are spectacularly exposed (Skipp et al., 2009). East of Arco, these limestones are painted with high-school graduation class years (Fig. 22). Continue through Butte City and then turn east at fork with Idaho 22 toward Idaho Falls on U.S. Highway 26. Stop 2-8 Big Lost River Rest Area on Highway 26 (12T 337790E; 4823650N; 1536 m; 4982 ft) Here the Big Lost River is normally dry. When it flows, water goes northward into the Big Lost Trough, a volcanically silled playa system that has existed since the late Pliocene. From the rest stop to the south, East, Middle and Big Southern Butte (see Figs. 4 and 10) mark the southwestern end of the Axial Volcanic Zone, a large volcanic highland that divides the Snake River from the streams that drain the mountain ranges north of the eastern Snake River Plain. East Butte is ~600,000 years old, Middle Butte is capped by tilted 1.0 Ma basalts and is younger than 1.0 Ma, and Big Southern Butte is ~300,000 years old (Kuntz et al., 1994). Cross section Figure 9, modified from Champion et al., (2011) extends southwest-northeast through this area. Far to the west, past the Craters of the Moon volcanic center which forms the skyline (Kuntz et al., 2007), two deep wells at Wendell and Mountain Home Air Force Base were sampled for detrital zircon analysis. Those analyses confirm that the Big Lost River flowed westward into the Snake River before the development of the Arco–Big Southern Butte Volcanic Rift Zone and Craters of the Moon volcanic center. ACKNOWLEDGMENTS The detrital-zircon geochronology summarized here was supported by National Science Foundation grants EAR-01-25756, 05-10980, and 08-38476. We thank reviewers Jennifer Pierce and Jeff Lee for many constructive improvements. The authors thank the Snake River Section of the Society for Mining, Metallurgy & Exploration for supporting publication of color figures. REFERENCES CITED Barton, I., 2007, Calculating the peak discharge of the outburst flood from Glacial Lake East Fork in the Big Lost River valley, east-central Idaho [M.S. thesis]: Bethlehem, Pennsylvania, Lehigh University.

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Beranek, L.P., Link, P.K., and Fanning, C.M., 2006, Miocene to Holocene landscape evolution of the western Snake River Plain region, Idaho: Using the SHRIMP detrital zircon provenance record to track eastward migration of the Yellowstone hotspot: Geological Society of America Bulletin, v. 118, p. 1027–1050, doi:10.1130/B25896.1. Bonnichsen, B., and Godchaux, M.M., 2002, Late Miocene, Pliocene, and Pleistocene geology of southwestern Idaho with emphasis on basalts in the Bruneau-Jarbidge, Twin Falls, and western Snake River Plain regions, in Bonnichsen, B., White, C.M., and McCurry. M., eds., Tectonic and magmatic evolution of the Snake River Plain volcanic province: Idaho Geological Survey Bulletin 30, p. 233–312. Champion, D.E., Hodges, M.K.V., Davis, L.C., and Lanphere, M.A., 2011, Paleomagnetic correlation of surface and subsurface basaltic lava flows and flow groups in the southern part of the Idaho National Laboratory, Idaho, with paleomagnetic data tables for drill cores at the Idaho National Laboratory: U.S. Geological Survey Scientific Investigations Report SIR (in press). Christiansen, R.L., 2001, The Quaternary and Pliocene Yellowstone Plateau Volcanic Field of Wyoming, Idaho, and Montana: U.S. Geological Survey Professional Paper 729-G, 145 p. Crone, A.J., 1987, Surface faulting and groundwater eruptions associated with the 1983 Borah Peak earthquake, in Link, P.K., and Hackett, W.R., eds., Guidebook to the geology of central and southern Idaho: Idaho Geological Survey Bulletin 27, p. 227–232. Diedesch, T., and Rodgers, D.W., 2010, Kinematic history of the Wildhorse detachment fault, Pioneer Mountains, south-central Idaho: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 197. Dover, J.H., 1981, Geology of the Boulder-Pioneer Wilderness Study Area, Blaine and Custer Counties, Idaho: Mineral resources of the BoulderPioneer Wilderness Study Area, Blaine and Custer Counties, Idaho: U.S. Geological Survey Bulletin 1497, p. 1–75. Dover, J.H., 1983, Geologic map and sections of the central Pioneer Mountains, Blaine and Custer Counties, central Idaho: U.S. Geological Survey Miscellaneous Investigations Series, Map I-1319, scale 1:48,000. Durk, K.M., 2007, Geochronology of part of the Wildhorse Gneiss Complex, Pioneer Mountains, Custer County, Idaho [unpublished senior thesis]: Pocatello, Idaho, Idaho State University Department of Geosciences, 38 p., http:// geology.isu.edu/dml/thesis/Durk_Kathleen_Senior_Thesis2007_ISU.pdf. Durk, K.M., Link, P.K., and Fanning, C.M., 2007, Neoproterozoic 695 Ma orthogneiss, Wildhorse Creek, Pioneer Mountains, south-central Idaho: New tie point in reconstruction of Rodinian rifting: Geological Society of America Abstracts with Programs, v. 36, no. 6, p. 613. Evenson, E.B., Cotter, J.F.P., and Clinch, J.M., 1982, Glaciation of the Pioneer Mountains: A proposed model for Idaho, in Bonnichsen, Bill, and Breckenridge, R.M., eds., Cenozoic geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 653–665. Evenson, E.B., Breckenridge, R.M., and Stephens, G.C., 1987, Field guides to the Quaternary geology of central Idaho, in Link, P.K., and Hackett, W.R., eds., Guidebook to the geology of central and southern Idaho: Idaho Geological Survey Bulletin 27, p. 201–244. Geslin, J.K., Link, P.K., and Fanning, C.M., 1999, High-precision provenance determination using detrital-zircon ages and petrography of Quaternary sands on the eastern Snake River Plain, Idaho: Geology, v. 27, p. 295– 298, doi:10.1130/0091-7613(1999)027<0295:HPPDUD>2.3.CO;2. Geslin, J.K., Link, P.K., Riesterer, J.W., Kuntz, M.A., and Fanning, C.M., 2002, Pliocene and Quaternary stratigraphic architecture and drainage systems of the Big Lost Trough, northeastern Snake River Plain, Idaho, in Link, P.K., and Mink, L.L., eds., Geology, hydrogeology and environmental remediation, Idaho National Engineering and Environmental Laboratory, eastern Snake River Plain, Idaho: Geological Society of America Special Paper 353, p. 11–26. Grader, G.W., and Dehler, C.M., 1999, Devonian stratigraphy in east-central Idaho: New perspectives from the Lemhi Range and Bayhorse Area, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the geology of eastern Idaho: Idaho Museum of Natural History, p. 31–36. Hart, W.K., and Brueseke, M.E., 1999, Analysis and Dating of Volcanic Horizons from Hagerman Fossil Beds National Monument and a Revised Interpretation of Eastern Glenns Ferry Formation Chronostratigraphy: A Report of Work Accomplished and Scientific Results Order No. 143PX9608-97-003, consolidated for electronic version, March 14, 2002, prepared for Hagerman Fossil Beds National Monument, P.O. Box 570, 221 N. State Street, Hagerman, Idaho 83332, USA.

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Hayden, F.V., 1883, Twelfth Annual Report of the U.S. Geological and Geographical Survey of the Territories—Wyoming and Idaho, for the year 1878: Washington D.C., Government Printing Office. Hodges, M.K.V., Link, P.K., and Fanning, C.M., 2009, Drainage disruption by basaltic lava fields along the Snake River Plain hotspot track: Detrital zircon evidence from drillcore constrains the course of the Pliocene Big Lost River: Journal of Volcanology and Geothermal Research, v. 188, p. 237–249, doi:10.1016/j.jvolgeores.2009.08.019. Ingersoll, R.V., Kretchmer, A.G., and Valles, P.K., 1993, The effect of sampling scale on actualistic sandstone petrofacies: Sedimentology, v. 40, p. 937– 953, doi:10.1111/j.1365-3091.1993.tb01370.x. Isaacson, P.E., and Dorobek, S.L., 1988, Regional significance and interpretation of a coral-stromatoporoid carbonate buildup succession, Jefferson Formation (Upper Devonian), east-central Idaho, in McMillan, N.J., Embry, A.F., and Glass, D.J., eds., Devonian of the World: Canadian Society of Petroleum Geologists, v. ii, p. 581–590. Isaacson, P.E., McFaddan, M.D., Measures, E.A., and Dorobek, S.L., 1988, Coral-stromatoporoid carbonate buildup succession, Jefferson Formation (Late Devonian), central Idaho, U.S.A., in Geldsetzer, H.H.J, James, N.P., and Tebbutt, G.E., eds., Reefs: Canada and adjacent area: Canadian Society of Petroleum Geologists Memoir 13, p. 471–477. Janecke, S.U. and Wilson, E., 1992, Geologic map of the Borah Peak, Burnt Creek, Elkhorn Creek, and Leatherman Peak 7.5-minute quadrangles, Custer and Lemhi Counties, Idaho: Idaho Geological Survey Map T-92-5, scale 1:24,000. Kenworthy, M.K., Pierce, J.L., and Rittenour, T., 2010, OSL chronology for alluvial fans of the Lost River Range, Idaho: Large scale deposition during OIS 3 and 4: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 74. Kuntz, M.A., Skipp, B.A., Lanphere, M.A., Scott, W.E., Pierce, K.L., Dalrymple, G.B., Champion, D.E., et al., 1994, Geologic map of the Idaho National Engineering Laboratory and adjoining areas, eastern Idaho. I-2330, Scale: 1:100,000. Kuntz, M.A., Skipp, B., Champion, D.E., Gans, P.B., Van Sistine, D.P., and Snyders, S.R., 2007, Geologic map of the Craters of the Moon 30′ × 60′ quadrangle, Idaho: U.S. Geological Survey Scientific Investigations Map 2969, 64-p. pamphlet, 1 plate, scale 1:100,000. Lewis, R.E., and Stone, M.A., 1988, Geohydrologic data from a 4403-foot geothermal test hole, Mountain Home Air Force Base, Elmore County, Idaho: U.S. Geological Survey Open-File Report 88-166, 30 p. Link, P.K., and Janecke, S.U., 1999, Geology of east-central Idaho: Geologic roadlogs for the Big and Little Lost River, Lemhi, and Salmon River Valleys, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the Geology of Eastern Idaho: Idaho Museum of Natural History, Pocatello, Idaho, p. 295–334. Link, P.K., Mahoney, J.B., Batatian, L.D., Bruner, D.J., and Williams, F., 1995, Stratigraphic setting of sediment-hosted mineral deposits in the eastern part of the Hailey 1° × 2° quadrangle, and part of the southern part of the Challis 1° × 2° quadrangle, south-central Idaho: U.S. Geological Survey Bulletin 2064-C, p. C1–C33. Link, P.K., Warren, I., Preacher, J.M., and Skipp, B., 1996, Stratigraphic analysis and interpretation of the Copper Basin Group, McGowan Creek Formation and White Knob Limestone, south-central Idaho, 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. 117–144. Link, P.K., McDonald, H.G., Fanning, C.M., and Godfrey, A.E., 2002, Detrital zircon evidence for Pleistocene drainage reversal at Hagerman Fossil Beds National Monument, central Snake River Plain, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 105–119. Link, P.K., Fanning, C.M., and Beranek, L.P., 2005, Reliability and longitudinal change of detrital-zircon age spectra in the Snake River system, Idaho and Wyoming: An example of reproducing the bumpy barcode: Sedimentary Geology, v. 182, p. 101–142, doi:10.1016/j.sedgeo.2005.07.012. Link, P.K., Fanning, C.M., Lund, K.I., and Aleinikoff, J.N., 2007, Detrital zircons, correlation and provenance of Mesoproterozoic Belt Supergroup and correlative strata of east-central Idaho and southwest Montana, in Link, P.K., and Lewis, R.S., eds., SEPM Special Publication 86, Proterozoic geology of western North America and Siberia, p. 101–128.

Link, P.K., Cameron, A., and Durk, A.K., 2010, Tectonostratigraphy of the Wildhorse gneiss complex, Pioneer Mountains: Neoarchean orthogneiss and overlying Mesoproterozoic Lemhi Group overlie Albion Range Proterozoic quartzite: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 415. Lund, K., 2008, Geometry of the Neoproterozoic and Paleozoic rift margin of western Laurentia: Implications for mineral deposit settings: Geosphere, v. 4, no. 2, p. 429–444, doi:10.1130/GES00121.1. Lund, K., Aleinikoff, J.N., Evans, K.V., duBray, E.A., Dewitt, E.H., and Unruh, D.M., 2010, SHRIMP U-Pb dating of recurrent Cryogenian and Late Cambrian–Early Ordovician alkalic magmatism in central Idaho: Implications for Rodinian rift tectonics: Geological Society of America Bulletin, v. 122, p. 430–453. McQuarrie, N., and Rodgers, D.W., 1998, Subsidence of a volcanic basin by flexure and lower crustal flow: The eastern Snake River Plain, Idaho: Tectonics, v. 17, p. 203–220, doi:10.1029/97TC03762. Michalek, M., 2009, Age and amount of crustal flexure in the Lake Hills, southcentral Idaho, and implications for the subsidence of the Eastern Snake River Plain [M.S. thesis]: Idaho State University, Pocatello, 111 p. Morgan, L.A., 1992, Stratigraphic relations and paleomagnetic and geochemical correlations of major ignimbrites of the eastern Snake River Plain, eastern Idaho and western Wyoming, in Link, P.K., Kuntz, M.A., and Platt, L.B., eds., Regional geology of eastern Idaho and western Wyoming: Geological Society of America Memoir 179, p. 215–227. Morgan, L.A., and McIntosh, W.C., 2005, Timing and development of the Heise volcanic field, Snake River Plain, Idaho, western USA: Geological Society of America Bulletin, v. 117, p. 288–306; doi: 10.1130/B25519.1. Ore, H.T., 1999, Topographic and geomorphic development of southeastern Idaho, segments from an essay, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the geology of eastern Idaho: Pocatello, Idaho, Idaho Museum of Natural History, p. 254–255. Patterson, S.J., 2006, Sedimentology and geomorphology of Quaternary alluvial fans with implications to growth strata, Lost River Range, Idaho [M.S. thesis]: Bozeman, Montana, Montana State University, 161 p. Peng, X., and Humpheys, E.D. 1998, Crustal velocity structure across the eastern Snake River Plain and the Yellowstone swell: Journal of Geophysical Research, v. 103, p. 7171–7186. Pierce, K.L., and Morgan, L.A., 1992, The track of the Yellowstone hot spot: Volcanism, faulting, and uplift, in Link, P.K., Kuntz, M.A., and Platt, L.B., eds., Regional geology of eastern Idaho and western Wyoming: Geological Society of America Memoir 179, p. 1–53. Pierce, K.L., and Morgan, L.A., 2009, Is the track of the Yellowstone hotspot driven by a deep mantle plume?—review of volcanism, faulting, and uplift in light of new data: Journal of Volcanology and Geothermal Research, v. 188, p. 1–25, doi:10.1016/j.jvolgeores.2009.07.009. Pierce, K.L., and Scott, W.E., 1982, Pleistocene episodes of alluvial-gravel deposition, southeastern Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 685–702. Rathburn, S.L., 1993, Pleistocene cataclysmic flooding along the Big Lost River, east-central Idaho: Geomorphology, v. 8, p. 305–319, doi:10.1016/0169 -555X(93)90026-X. Rodgers, D.W., Link, P.K., and Huerta, A.D., 1995, Structural framework of mineral deposits hosted by Paleozoic rocks in the northeastern part of the Hailey 1° × 2° quadrangle, south-central Idaho: U.S. Geological Survey Bulletin 2064-B, p. B1–B18. Rodgers, D.W., Ore, T.H., Bobo, R.T., McQuarrie, N., and Zentner, N., 2002, Extension and subsidence of the eastern Snake River Plain, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 121–155. Ross, C.P., 1947, Geology of the Borah Peak quadrangle, Idaho: Geological Society of America Bulletin, v. 58, no. 12, p. 1085–1160, doi:10.1130/0016-7606(1947)58[1085:GOTBPQ]2.0.CO;2. Sadler, J., and Link, P.K., 1996, The Tuana Gravel; early Pleistocene response to longitudinal drainage of a late-stage rift basin, western Snake River plain, Idaho: Northwest Geology, v. 26, p. 46–64. Scott, W.E., Pierce, K.L., and Hait, M.H., Jr., 1985, Quaternary tectonic setting of the 1983 Borah Peak earthquake, central Idaho: in Proceedings of Workshop XXVIII on the Borah Peak, Idaho, Earthquake: U.S. Geological Survey Open-File Report 85-290, v. A, p. 1–16.

The Neogene drainage history of south-central Idaho Sheedy, V.E., 2003, Surficial geology and sediment provenance: A question of stream capture in the Big Lost River and Antelope Flat regions, Custer and Butte Counties, Idaho [M.S. thesis]: Pocatello, Idaho State University, 140 p. Skipp, B., Snider, L.A., Janecke, S.U., and Kuntz, M.A., 2009, Geologic Map of the Arco 30′ × 60′ quadrangle, Idaho: Idaho Geological Survey Map 47, scale 1:100,000. Stewart, E.D., Link, P.K., Fanning, C.M., Frost, C.D., and McCurry, M., 2010, Paleogeographic implications of non–North American sediment in the Mesoproterozoic upper Belt Supergroup and Lemhi Group, Idaho and Montana, USA: Geology, v. 38, p. 927–930, doi:10.1130/G31194.1. Stroup, C.N., Link, P.K., and Fanning, C.M., 2008a, Provenance of Late Miocene fluvial strata of the Sixmile Creek Formation, southwest Montana: Evidence from detrital zircon: Northwest Geology, v. 37, p. 69–84. Stroup, C.N., Link, P.K., Janecke, S.U., Fanning, C.M., Yaxley, G.M., and Beranek, L.P., 2008b, Eocene to Oligocene provenance and drainage in extensional basins of southwest Montana and east-central Idaho: Evidence from detrital zircon populations in the Renova Formation and equivalent strata, in Spencer, J.E., and Titley, S.R., eds., Circum-Pacific tectonics, geologic evolution, and ore deposits: Arizona Geological Survey Digest 22, p. 529–546. Sutfin, N.A., Sharp, W., Pierce, J.L., and Kenworthy, M.K., 2010, Estimating the age of terminal fluvial deposition on Quaternary fans of the Lost River basin, northern Rocky Mountains, USA: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 470. Tauxe, L., Luskin, C., Selkin, P., Gans, P., and Calvert, A., 2004, Paleomagnetic results from the Snake River Plain: Contribution to the time-averaged field

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global database: Geochemistry Geophysics Geosystems, v. 5, Q08H13, doi:10.1029/2003GC000661. Thackray, G.D., Rodgers, D.W., and Drabick, A., 2011, this volume, Neogene drainage development of Marsh and Portneuf valleys, eastern Idaho in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, doi: 10.1130/2011.0021(04). Vogl, J.J., Carmenate, A., Min, K., Rodgers, D.W., and Fanning, C.M., 2010, Extensional and cooling history of the Pioneer core complex, south-central Idaho: Long-lived episodic extension and implications for regional patterns of extension: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 595. Whitehead, R.L., 1992, Geohydrologic framework of the Snake River Plain aquifer system, Idaho and eastern Oregon: U.S. Geological Survey Professional Paper 1408-B, 32 p. Williams, I.S., 1998, U-Th-Pb Geochronology by Ion Microprobe, in McKibben, M.A., Shanks W.C., III, and Ridley, W.I., eds., Applications of microanalytical techniques to understanding mineralizing processes: Reviews in Economic Geology, v. 7, p. 1–35. Wust, S.L., 1986, Extensional deformation with northwest vergence, Pioneer core complex, central Idaho: Geology, v. 14, p. 712–714, doi:10.1130/0091-7613(1986)14<712:EDWNVP>2.0.CO;2.

MANUSCRIPT ACCEPTED BY THE SOCIETY 9 MARCH 2011

Printed in the USA

The Geological Society of America Field Guide 21 2011

Paleontology and stratigraphy of middle Eocene rock units in the Bridger and Uinta Basins, Wyoming and Utah Paul C. Murphey* Department of Paleontology, San Diego Natural History Museum, 1788 El Prado, Balboa Park, San Diego, California 92101, USA K.E. Beth Townsend* Arizona College of Osteopathic Medicine, Midwestern University, 19555 N. 59th Avenue, Glendale, Arizona 85308, USA Anthony R. Friscia* Department of Integrative Biology and Physiology, University of California, Los Angeles, 621 Charles E. Young Drive So., Los Angeles, California 90095-1606, USA Emmett Evanoff* Department of Earth Sciences, University of Northern Colorado, Greeley, Colorado 80639, USA

A grand scene burst open us. Fifteen hundred feet below us lay the beds of another great Tertiary lake. We stood upon the brink of a vast basin so desolate, wild, and broken, so lifeless and silent, that it seemed like the ruins of the world. —Charles Betts’ description of the first glimpse of the Uinta basin in his account of the 1870 Yale College Expedition led by paleontologist O.C. Marsh, published in Harper’s magazine, 1871, v. 43, p. 66.

ABSTRACT The Bridger Formation is located in the Green River basin in southwest Wyoming, and the Uinta and Duchesne River formations are located in the Uinta basin in Utah. These three rock units and their diverse fossil assemblages have great scientific importance and are also of historic interest to vertebrate paleontologists. Notably, they are also the stratotypes for the three middle Eocene North American Land Mammal “Ages,” the Bridgerian, Uintan, and Duchesnean, from oldest to youngest. The fossils and sediments of these formations provide a critically important record of biotic, environmental, and climatic history spanning ~10 million years (49–39 Ma). This article features a detailed field excursion through portions of the Green River and Uinta basins that focuses on locations of geologic, paleontologic, and historical interest. In support of the field excursion, we also provide a review of current knowledge of these formations with emphasis on lithostratigraphy, biochronology, depositional and paleoenvironmental history, and the history of scientific exploration. *[email protected]; [email protected]; [email protected]; [email protected]. Murphey, P.C., Townsend, K.E.B., Friscia, A.R., and Evanoff, E., 2011, Paleontology and stratigraphy of middle Eocene rock units in the Bridger and Uinta Basins, Wyoming and Utah, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 125–166, doi:10.1130/2011.0021(06). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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INTRODUCTION AND STRUCTURAL SETTING Situated to the north and south of the Uinta Mountains in Wyoming and Utah, respectively, the Bridger and Uinta basins have great scientific importance and are of historic interest to vertebrate paleontologists. The rock units and fossils of the Bridger basin (actually part of the southern Green River basin) and the Uinta basin have been the focus of paleontological investigations for the past 140 years. Perhaps the most familiar of the rock units within the Green River and Uinta basins is the Green River Formation because of its economic importance and exquisitely preserved vertebrate, invertebrate and plant fossils. However, despite the geological and paleontological importance of this world renowned lacustrine rock unit, this field excursion is focused on three closely related stratigraphically adjacent and overlying fluvial rock units that are best known for their assemblages of middle Eocene vertebrate fossils. The Bridger, Uinta, and Duchesne River formations are the stratotypes for the Bridgerian, Uintan and Duchesnean North American Land Mammal “Ages” (NALMAs) (Gunnell et al., 2009; Wood et al., 1941). The fossils and sediments of these formations provide a critically important record of biotic, environmental, and climatic history spanning ~10 million years (49–39 Ma).

The greater Green River basin occupies 32,187 km2 of southwestern Wyoming and northwestern Colorado (Roehler, 1992a). Structurally, it is a large asymmetrical syncline with mostly gently dipping flanks (3° to 5°) with steeper dips along the southern margin of the basin, and an approximately north-south axis (Koenig, 1960; Roehler, 1992a). The greater Green River basin is divided into four smaller basins by three intrabasin arches. The largest of these arches, the north-south–trending Rock Springs uplift, divides the basin into roughly equal halves, with the Green River basin to the west, and the Great Divide, Sand Wash, and Washakie basins to the east. The Bridger basin is located within the southern part of the Green River basin. The term Bridger basin (Hayden, 1871) traditionally refers to an area located north of the Uinta Mountains and south of the Blacks Fork of the Green River in Uinta and Sweetwater counties, Wyoming, and is a physiographic, not a structural basin (Fig. 1). The Uinta basin occupies 10,943 km2 of northeastern Utah. Structurally, it is an asymmetrical, elongate east-west–trending synclinal basin bounded by the Uinta Mountains to the north, the Douglas Creek Arch and Roan Plateau to the east, the Book Cliffs/Tavaputs Plateau to the south, and the Wasatch Range to the west (Fig. 2). It was formed in the latest Cretaceous and Paleocene during the Laramide uplift of the Uinta Mountains.

Figure 1. Index map of the greater Green River basin showing the approximate location of the Bridger basin (type area of the Bridger Formation), major structural features, and surrounding uplifts (modified from Murphey and Evanoff, 2007).

Middle Eocene rock units in the Bridger and Uinta Basins The Uinta basin is closely related structurally and sedimentologically to the Piceance Creek basin in northwestern Colorado. The Uinta and Piceance Creek basins are separated by the Douglas Creek arch, a broad north-south–trending anticline that separated the two sedimentary basins until the early-middle Eocene, when the two basins coalesced across the top of the arch to form one large sedimentary basin (Moncure and Surdam, 1980; Johnson, 1985, 1989). Like the Uinta basin, the Piceance Creek basin is highly asymmetrical (Johnson, 1985). Early Cenozoic strata in the Uinta basin dip gently from all directions to the northern margin of the basin, where the strata are sharply upturned and faulted along the southern flank of the Uinta Mountains uplift (Johnson, 1985). The greater Green River, Uinta, and Piceance Creek basins began forming during the Laramide orogeny, a period of tectonism in western North America that was initiated during the late Cretaceous and continued for ~30 million years until the late Eocene. In addition to the uplifting of surrounding mountain ranges, Laramide tectonism resulted in rapid subsidence in basin depositional centers, and lacustrine and fluvial deposition in these intermontane basins was mostly continuous. Lacustrine

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deposition was characterized by a complex history of expansions and contractions in response to basin subsidence, climatic conditions, and volcanic activity (Murphey, 2001; Murphey and Evanoff, 2007; Roehler, 1992b). PART I. BRIDGER BASIN FIELD TRIP With its abundant and diverse vertebrate fossils and extensive exposures, the Bridger Formation provides an excellent opportunity to study middle Eocene continental environments of North America. The dramatic and picturesque Bridger badlands are an 842 m (2763 ft) thick sequence dominated by green-brown and red mudstone and claystone, with interbedded scattered ribbon and sheet sandstone, widespread beds of micritic, sparry, and silicified limestone, and thin but widespread beds of ash-fall tuff (Evanoff et al., 1998; Murphey and Evanoff, 2007). This field trip offers participants the opportunity to examine paleontologically significant strata of the Bridger Formation in the southern Green River basin. The following sections of the field trip guide provide a summary of the Cenozoic geologic history of the Green River basin, as well as the history of

Figure 2. Index map of the Uinta Basin showing adjacent structural and physiographic features. The gray region represents the areal extent of Tertiary deposits in both the Uinta basin, Utah, and Piceance Creek basin (not labeled), western Colorado. The Douglas Creek Arch separates the Uinta basin from the Piceance Creek basin.

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investigations, stratigraphy, depositional and paleoenvironmental history, and fossils of the Bridger Formation. This is followed by a detailed road log. Paleogene Geologic History of the Green River Basin, Wyoming The greater Green River basin was filled with Paleocene and Eocene fluvial and lacustrine sediments, and, during the Eocene, sedimentation appears to have been continuous in most of the basin. The oldest Cenozoic rock units in the greater Green River basin, the Paleocene Fort Union Formation and the early Eocene Wasatch Formation, are exposed mostly along its eastern and western flanks. During the Paleocene and earliest Eocene, deposition in the greater Green River basin was predominantly fluvial, with epiclastic sediments accumulating in river drainages and on adjacent floodplains. The onset of lacustrine deposition associated with the Green River lake system may have commenced as early as the late Paleocene (Grande and Buchheim, 1994). Lake sediments accumulated on broad floodplains of low topographic relief, and the lake waters expanded and contracted numerous times over the next approximately five million years in response to climatic changes, tectonic influences, and episodic volcanic activity. Occupying the center of the basin in the shape of a large, irregular lens (Bradley, 1964; Roehler, 1992b, 1993), the Green River Formation is the result of at least five million years of lacustrine deposition lasting from ca. 53.5 to 48.5 Ma (Smith et al., 2003), although lacustrine deposition may have persisted later in the southernmost part of the basin along the Uinta Mountain front (Murphey and Evanoff, 2007). The Green River Formation was deposited in a vast ancient lake system that existed from the late Paleocene to the middle Eocene in what is now Colorado, Utah, and Wyoming. The smallest and oldest of these lakes, Fossil Lake, was deposited in Fossil basin, which is located in the Wyoming overthrust belt just to the west of the Green River basin in southwestern Wyoming. Lake Gosiute was deposited in the greater Green River basin, which includes the Green River and Washakie basins in southwestern Wyoming, and the Sand Wash basin in northwestern Colorado. Fossil Lake and Lake Gosiute may never have been physically connected (Surdam and Stanley, 1980). Lake Uinta was deposited in the Uinta basin in northeastern Utah and the Piceance Creek basin in northwestern Colorado. Lithologically, the Green River Formation in the greater Green River basin is a complex sequence of limestone, shale, and sandstone beds with a maximum thickness of ~838 m (2750 ft) (Roehler, 1993). It was deposited lateral to and above the predominantly fluvial Wasatch Formation, and lateral to and below the fluvial and lacustrine Bridger and Washakie formations. The Laney Member is the uppermost member of the Green River Formation in Wyoming and represents the final expansion of Lake Gosiute. Most volcaniclastic sediments deposited in the Green River basin during the middle Eocene were apparently transported from

the Absaroka Volcanic Field in what is now northwestern Wyoming. These sediments were washed into the basin in rivers and streams. Some volcaniclastic sediments were transported into the basin via eolian processes and deposited as ash fall in lakes and on floodplains. A large influx of fluvially transported volcaniclastic sediment is believed to have led to the final middle Eocene filling of Lake Gosiute (Mauger, 1977; Murphey, 2001; Murphey and Evanoff, 2007; Surdam and Stanley, 1979). Mauger (1977) and Surdam and Stanley (1979) estimated that Lake Gosiute was ultimately extinguished by ca. 44 Ma. The Bridger, Green River, and Washakie formations are locally and unconformably overlain by the Oligocene Bishop Conglomerate and the middle-to-late–Miocene Browns Park Formation. Since the Eocene, the greater Green River basin has been modified by erosion, regional uplift, and normal faulting, but the basic structure of the basin remains the same as it was during deposition of the Wasatch, Green River, Washakie, and Bridger formations. History of Paleontological Investigations in the Bridger Formation John Colter, who traveled to the headwaters of the Green River in 1807, was probably among the first non–Native Americans to visit the Green River basin (Chadey, 1973). Hundreds of subsequent trappers and explorers traversed the basin during the first half of the nineteenth century, and a number of records of these early explorations make reference to fossils and coal (Roehler, 1992a). The earliest scientific observations on the geology of the Green River basin were made by Army Lt. John C. Fremont. After entering the basin through South Pass at the southern end of the Wind River Mountain, Fremont (1845) described varicolored rocks (now known as Eocene-age Wasatch Formation) along the Big Sandy and New Fork rivers. He also collected fossil shells from near Cumberland Gap (Veatch, 1907). The earliest vertebrate fossils reported from the Green River basin were fishes discovered in the Green River Formation. In 1856, Dr. John Evans collected a specimen of a fossil fish from an unknown Green River Formation locality west of Green River City. He sent this specimen to paleontologist Joseph Leidy in Philadelphia for study, and Leidy named it Clupea humilis (later renamed Knightia humilis) (West, 1990). Hayden (1871) described the discovery of a locality he referred to as the “petrified fish cut” along the main line of the Union Pacific Railroad ~2 miles west of Green River. Employees of the railroad had initially discovered the locality and later turned many specimens over to Hayden. Paleontologist Edward Drinker Cope described the fish fossils from the petrified fish cut in Hayden’s (1871) expedition report. The initial discovery of mammalian fossils in the Green River basin was probably made by a long-time local resident. Trapper Jack Robinson (also called Robertson) found what he described as a “petrified grizzly bear” sometime in the late 1860s in what is now called the Bridger Formation but which had initially been

Middle Eocene rock units in the Bridger and Uinta Basins named the “Bridger Group” by Ferdinand V. Hayden in 1869. This story was related to Joseph Leidy by Judge William Carter of Fort Bridger as an explanation for the name “Grizzly Buttes,” an area 10–15 miles southeast of Fort Bridger where fossils were particularly common (the name Grizzly Buttes has since disappeared from the local geographic vocabulary). Several government geological and topographical surveys with specific but overlapping territories were operating in the southern Green River basin between 1867 and 1879. Hayden and his party collected along the Henrys Fork valley and further north in the vicinity of Church Buttes in 1870 as part of the 1867– 1878 U.S. Geological and Geographical Survey of the Territories (Hayden, 1873). Fossils collected by Hayden’s group were sent to Joseph Leidy in Philadelphia for study and were described in his 1873 monograph on fossil vertebrates. Later paleontological studies for the Hayden Survey were carried out by E.D. Cope. Under the direction of John Wesley Powell, the U.S. Geological and Geographical Survey of the Territories, Second Division (1875–1876), worked along the Henrys Fork River in 1869, and in a corridor 10–20 miles wide on either side of the Green River in 1871 (Powell, 1876). The U.S. Geological Survey of the Fortieth Parallel (1867–1872), directed by Clarence King, worked in the Green River basin in 1871 and 1872. The fossils collected by the King Survey were sent to Othniel Charles Marsh for description. Most of the fossils collected during these surveys were discovered in the Bridger Formation. Many of the early scientific expeditions to the Green River basin were based out of Fort Bridger that was originally set up as a trading post in 1843 by trapper and guide Jim Bridger and his partner Louis Vasquez. The fort became an army post after the 1857 Mormon War. Judge Carter and Dr. J. Van A. Carter, later residents of Fort Bridger, maintained an active correspondence with Joseph Leidy in Philadelphia during the late 1860s and early 1870s. This correspondence included mailing fossils to Leidy, which were described in subsequent publications (Leidy, 1869, 1871, 1872a, 1872b, 1873). Leidy, who is often regarded as the father of North American vertebrate paleontology (Lanham, 1973), named the first Bridger Formation fossil to be formally described, the omomyid primate Omomys carteri, after Dr. Carter (Leidy, 1869). Omomys carteri was also the first-described fossil primate from North America. Early reports of fossils from the Green River basin did not go unnoticed by rival paleontologists O.C. Marsh and E.D. Cope. The incidents that set the stage for the long and bitter conflict between these two men began in the Green River basin while they were prospecting in the Bridger Formation in 1872. Sometimes referred to as the “bone wars,” the dispute between Marsh and Cope lasted for more than 30 years and included efforts by each man to destroy the scientific reputation and integrity of the other. This conflict soured Leidy’s interest in paleontology and led to his eventual abandonment of the discipline after 1872. Professor Marsh was the first professional paleontologist to collect fossils from the Bridger Formation; he brought crews with him from Yale College for four consecutive summers from 1870

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to 1873. Leidy’s only excursion to the West took place in 1872, when he visited the Bridger badlands guided by the Carter brothers of Fort Bridger. Cope’s only visit to the Bridger badlands occurred in 1872, while Cope was attached to the Hayden survey as the paleontologist. This visit infuriated Marsh, who, at the time, considered the Green River basin and Bridger Formation his exclusive fossil-collecting territory. By the late 1870s, Cope and Marsh had left the Green River basin for good, although both men independently and at different times retained the services of paid fossil collector Sam Smith (West, 1990). Other early fossil collectors who visited the Green River basin in 1877 and 1878 included Henry Fairfield Osborn, William Berryman Scott, and Francis Speir for Princeton University. Scott returned to the area with Speir in 1886. Jacob Wortman and James W. Gidley collected for the American Museum of Natural History (AMNH) in 1893. The early fossil-collecting expeditions to the Green River basin resulted in large collections of fossils primarily from the Bridger Formation at the Philadelphia Academy of Natural Sciences (Leidy), Yale University (Marsh), the AMNH (which purchased Cope’s collection just before the turn of the century), and Princeton University (Osborn, Scott, and Speir). Unfortunately, these early collectors paid little attention to the stratigraphic provenance of the fossils they collected. Their collections do, nevertheless, contain the holotypes of most presently recognized Bridgerian mammal taxa. In 1902, H.F. Osborn, who was then the U.S. Geological Survey paleontologist, initiated the first program of stratigraphic fossil collection to take place in the Green River basin and one of the first in North America. Osborn charged Walter Granger and William Diller Matthew of the AMNH with the task of carrying out the study. Matthew was also directed to find a uintathere to display at the AMNH. The AMNH party, led by Granger, worked in the Bridger basin from 1902 to 1906 (Matthew, 1909). The second halves of the 1903 and 1905 field seasons were devoted to mapping and describing the stratigraphy of the Bridger Formation, while the remainder of the time was spent searching the badlands for fossils. The efforts of the AMNH parties over these four years resulted in an excellent fossil collection that was, for its time, very well documented stratigraphically. These AMNH expeditions also resulted in the first paper to be published on the geology of the Bridger Formation, which was authored by William J. Sinclair (1906), who had joined the AMNH field party for the summer of 1905. In Matthew’s classic 1909 monograph, The Carnivora and Insectivora of the Bridger Basin, Middle Eocene, the geology of the Bridger Formation was described briefly, and a system of stratigraphic subdivisions for the formation was introduced. These subdivisions, Bridger A–E, were based on areally extensive limestone beds, which Matthew called “white layers.” Following the early fossil-collecting expeditions of the nineteenth century and initial scientific field studies conducted by AMNH crews in the early twentieth century, the Bridger Formation in the Green River basin has remained the focus of almost continuous paleontologic inquiry because of its abundant and

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diverse vertebrate fossils, although Matthew’s (1909) original stratigraphy was only recently refined. West (1990) wrote an excellent historical summary of vertebrate paleontological work in the Green River basin from 1840 to 1910. H.F. Osborn (1929) devoted considerable discussion to the Bridger Formation and its fossils in his monograph, The Titanotheres of Ancient Wyoming, Dakota, and Nebraska. Horace Elmer Wood (1934) divided the Bridger Formation into two members. The Blacks Fork Member corresponds to Matthew’s Bridger A and B, and the Twin Buttes Member corresponds to Matthew’s Bridger C and D, with the Sage Creek White Layer marking their boundary. Contrary to rules of stratigraphic nomenclature, these members were defined on perceived faunal differences rather than lithologic differences. The informal usage of the terms “Blacksforkian” and “Twinbuttean” as land mammal subages derives from the names of the two Bridger members. Under the direction of J.W. Gidley, followed by C. Lewis Gazin, the Smithsonian Institution began an active collecting program in the Bridger Formation beginning in 1930. Gazin was active in the Green River basin from 1941 to 1968. This period of activity resulted in a relatively large and well-documented collection that was the subject of numerous publications by Gazin focused primarily on the systematic paleontology of Bridgerian mammal fossils (e.g., 1934, 1946, 1949, 1957, 1958, 1965, 1968, and 1976). Paul O. McGrew and Raymond Sullivan worked on the stratigraphy and paleontology of the Bridger A in the late 1960s and published the results of their work in 1970. Robert M. West began an active collecting program for the Milwaukee Public Museum in 1970 and worked in the basin until the late 1970s. West’s work, which also resulted in a large number of paleontological publications, included the use of screen-washing techniques to collect microvertebrates, a portion of the fauna that had not been previously well sampled. Like Wood (1934) and Koenig (1960), West (1976) noted difficulties with the correlation of Matthew’s white layers across the basin and suggested that a bipartite division of the Bridger into upper (Twin Buttes) and lower (Blacks Fork) members was most appropriate. West and Hutchison (1981) named Matthew’s Bridger E the Cedar Mountain Member, adding a third member to the Bridger Formation. Paleontological and geological studies of Tabernacle Butte, an isolated remnant of the Bridger Formation of late Bridgerian age with an important fossil fauna, were published by McGrew (1959), McKenna et al. (1962), and West and Atkins (1970). Evanoff et al. (1998), Murphey (2001), and Murphey and Evanoff (2007) significantly refined Matthew’s (1909) Bridger Formation stratigraphic scheme. Their work included the addition of newly described marker units; the establishment of new stratigraphic subdivisions and correlation of marker units across the southern part of the basin where the most complete stratigraphic sequence is exposed; descriptions of detailed stratigraphic sections measured through the Bridger B, C, D, and E; renaming of the Cedar Mountain Member to the Turtle Bluff Member in order to conform with the rules of stratigraphic nomenclature; stratigraphic positioning of more than 500 fossil localities; iso-

topic dating of four ash-fall tuffs; and geologic mapping of more than 600 mi2 of the southern Green River basin at the scale of 1:24,000. Geologic maps and publications relating to the Bridger Formation are available at http://www.rockymountainpaleontology.com/bridger. Stratigraphy and Depositional Environments of the Bridger Formation The Bridger Formation was named the “Bridger Group” by Hayden (1869). The first stratigraphic framework for the Bridger Formation was established by W.D. Matthew (1909) of the AMNH in the southern Green River basin where the formation is thickest and best exposed. Matthew’s (1909) stratigraphic subdivisions of the Bridger Formation were based primarily on five areally extensive limestone beds. These he named the Cottonwood, Sage Creek, Burnt Fork, Lonetree, and upper white layers, and some were used to subdivide the formation into five units: Bridger A, B, C, D, and E, from lowest to highest. Matthew’s intent was to make it possible to stratigraphically locate the numerous known fossil localities in the formation. Because they are the most fossiliferous, the Bridger B, C, and D were further divided into five subunits corresponding to basal, lower, middle, upper, and top levels (e.g., B1, B2, B3, B4, B5). Because Matthew (1909) did not define the upper and lower boundaries of these subunits with stratigraphic markers or measured sections, correlations between them and the later subdivisions proposed by Evanoff et al. (1998), Murphey (2001), and Murphey and Evanoff (2007) are uncertain. The history of stratigraphic nomenclature for the Bridger Formation is provided in Figure 3. In his 1909 monograph, Matthew (1909, p. 296) gave a brief description of his proposed five members and his white layers. “Horizon A” was 200 ft thick, composed primarily of calcareous shales alternating with tuffs, and with rare fossils. “Horizon B” was 450 ft thick, consisting of two benches separated by the Cottonwood white layer and containing abundant and varied fossils. He went on to note that the largest number of complete skeletons from the entire formation was found in the lower part of Horizon B (B2). “Horizon C” was 300 ft thick, “defined inferiorly” by the Sage Creek white layer, with the Burntfork white layer occurring at about its middle, and with abundant and varied fossils. He also noted that the Sage Creek white layer was the “heavy and persistent calcareous stratum” at Sage Creek Spring, thus designating a type locality where this unit had been previously described and illustrated, but not named, by Sinclair (1906). “Horizon D” was 350 ft thick, composed of harder gray and greenish-gray sandy and clayey tuffs, “defined inferiorly” by the Lonetree white layer, with the upper white layer ~75 ft from the top, and with abundant and varied fossils. “Horizon E” was 500 ft thick, composed of soft banded tuffs with heavy volcanic ash layers, with a high gypsum content and nearly barren of fossils. The total thickness of the Bridger reported by Matthew was 1800 ft. Despite the lithologic descriptions of the five horizons made by Matthew (1909), subsequent workers have not been able to subdivide the Bridger

Middle Eocene rock units in the Bridger and Uinta Basins Formation on the basis of lithologic differences (Bradley, 1964; Evanoff et al., 1998; Murphey, 2001, Murphey and Evanoff, 2007; Roehler, 1992a). Furthermore, with the exception of the Bridger B-C and D-E boundaries, Matthew’s subdivisions do not correspond to major faunal changes (Murphey, 2001; Murphey and Evanoff, 2007; Simpson, 1933; Wood, 1934). The Bridger Formation has been subdivided into three members. The Blacks Fork Member, or lower Bridger, is equivalent to Matthew’s Bridger A and B; the Twin Buttes Member, or upper Bridger, is equivalent to Matthew’s C and D; and the Turtle Bluff Member, also considered part of the upper Bridger, is equivalent to Matthew’s Bridger E. A detailed history of geologic and paleontologic investigations focusing on the Bridger Formation, and the history of stratigraphic nomenclature for this unit, are provided by Murphey and Evanoff (2007). Evanoff et al. (1998), Murphey (2001), and Murphey and Evanoff (2007) published the first major stratigraphic revision of the Bridger Formation since Matthew’s (1909) stratigraphy. The most recent stratigraphic subdivisions are based on widespread limestone beds, tuffs, and tuffaceous sheet sandstones which are used as marker units. Fifteen such units were described, and seven of these were considered major markers. These were used to subdivide the Bridger C and D (Twin Buttes Member) into lower, middle, and upper informal subdivisions (Fig. 3). Two

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additional markers were used to redefine the base and define the top of the Bridger E (Turtle Bluff Member). Four of Matthew’s original “white layers” were included in the stratigraphy of the Bridger C and D, and these were mapped and redescribed in detail. In conjunction with the latest stratigraphic revision, geologic mapping of ten 7.5 min quadrangles which cover the area encompassed by the upper Bridger Formation was completed, and these maps are available from the Wyoming State Geological Survey. Because many marker units are not continuously exposed or traceable across the entire basin (from Hickey Mountain, Sage Creek Mountain, and Cedar Mountain east to Twin Buttes and Black Mountain), a distance of ~40 miles, accurate correlation was made possible by using the mineralogically diagnostic Henrys Fork tuff as a datum. Rock accumulation rates, isotopic ages of ash-fall tuffs (Murphey et al., 1999), and fossils indicate that the 842 m (2763 ft) thick Bridger Formation was deposited over an ~3.5-millionyear interval from ca. 49.09 to 45.57 Ma, and that the faunal transition from the Bridgerian to the Uintan Land Mammal Age was under way by ca. 46 Ma as indicated by fossils collected from the Turtle Bluff Member (Evanoff et al., 1994; Gunnell et al., 2009; Murphey, 2001; Murphey and Evanoff, 2007; Robinson et al., 2004). Recognized depositional environments of the Bridger Formation include fluvial, lacustrine, playa lacustrine,

Figure 3. History of Bridger Formation stratigraphic nomenclature Bridger from 1869 until present. The correlation between Matthew’s (1909) subdivisions (1–5) and the lower, middle, and upper subdivisions of Murphey and Evanoff (2007) are uncertain.

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paludal, marginal mudflat, basin margin, and volcanic. Murphey and Evanoff (2007) concluded that an influx of fluvially transported volcaniclastic sediment to the Green River basin during middle Eocene time led to the filling of Lake Gosiute and the development of muddy floodplains of low topographic relief, which persisted for up to 85% of the time during which the upper Bridger was deposited. Occasional lapses in the flow of sediment to the basin permitted the development of shallow, mostly groundwater-fed lakes and ponds, which accumulated up to four times as slowly as floodplain deposits. These lapses decreased in frequency throughout deposition of the upper Bridger Formation. As indicated by fossil distribution and diversity, lakes and their margins provided favorable habitats for both aquatic and terrestrial organisms during deposition of the Bridger Formation. Middle Eocene Paleoenvironments of the Green River Basin Numerous studies based on paleontological and geological evidence have concluded that the Eocene-age rock units in the greater Green River basin were deposited in warm temperate, subtropical, and tropical climatic conditions (Roehler, 1993). Perhaps the most reliable information concerning paleoclimates comes from analysis of plant mega- and micro-fossils. According to Leopold and MacGinitie (1972), early Eocene floras (based on palynology of samples collected from the Niland Tongue of the Wasatch Formation and the Luman and Tipton tongues of the Green River Formation) suggest a humid subtropical to warm temperate climate with summer rainfall and only mild frost and with a mean annual temperature of 55 °F. Nichols (1987) concluded that the climate of the basin floor during deposition of the Niland Tongue was subtropical, without freezing temperatures. The earliest middle Eocene climates pertaining to the Cathedral Bluffs Tongue of the Wasatch Formation and the Wilkins Peak Member of the Green River Formation were interpreted as generally hot and dry (Leopold and MacGinitie, 1972). Climatic conditions in the early-middle Eocene during deposition of the lower part of the Laney Member of the Green River Formation were characterized as warm and humid with tropical affinities. Floras of the upper part of the Laney Member indicate a change to cooler, subhumid conditions (Leopold and MacGinitie, 1972). Both pollen and leaf data from the Washakie Formation indicate a dry but temperate climate (Leopold and MacGinitie, 1972). Roehler (1993) reported in a written communication that MacGinitie reinterpreted temperature and precipitation ranges on the basis of palynology of samples collected from the Washakie basin by Roehler (1992a). His reinterpretation estimated mean annual temperatures of 65 °F during the early Eocene, 63 °F during the earliest middle Eocene, and 62 °F during the middle Eocene. Average annual precipitation was estimated at more than 40 inches during the early Eocene, 25–35 inches during the earliest middle Eocene, and 15–20 inches in the middle Eocene. Sedimentological evidence of a more arid climate during the middle Eocene (transitional Uintan NALMA) includes massive beds of gypsum capping the Turtle Bluff Member of the Bridger

Formation (Murphey, 2001; Murphey and Evanoff, 2007). The shift from dominantly tropical forest environments to more-open, savanna-like conditions in the Eocene intermontane basins during late Bridgerian (early-middle Eocene) and Uintan (middle Eocene) times has also been studied by using ecological diversity analysis applied to mammalian faunas (Murphey and Townsend, 2005; Townsend, 2004). As indicated by fossil distribution and diversity, the Green River lakes and their forested margins provided highly favorable habitats and preservational environments for both aquatic and terrestrial organisms. Lake margin habitats, riparian corridors and adjacent floodplains were apparently vegetated during much of the time of Green River Formation deposition, as indicated by a paleoflora that includes a variety of trees and bushes such as palm, cinnamon, oak, maple, lilac, and hazel, as well as cattails and rushes. Insects of many varieties lived in the lakes and forests and are locally well preserved in lake sediments. A variety of terrestrial and aquatic mollusks (clams and snails) are also known to have inhabited the Green River lakes (Hanley, 1974). Crayfish, prawn, and ostracods inhabited the warm lake waters, as did a diversity of fish species, including relatives of the herring, perch, paddlefish, bowfin, gar, catfish, and stingray (Grande, 1984; Grande and Buchheim, 1994; McGrew and Casilliano, 1975). Frogs, crocodiles, and turtles were common residents of shallower proximal shoreline waters. A diversity of reptile species, including tortoise, lizards, and snakes, inhabited the forests surrounding Eocene lakes and ponds. Flamingos, hawks, rails, stone curlews, and other bird species frequented the forests, wetlands, and lakes (Murphey et al., 2001). The forests teemed with the primitive ancestors of many modern mammalian groups, including rodents, insectivores, bats, primates, perissodactyls (horse, rhinoceros, and tapir), and carnivores, as well as more bizarre archaic forms such as creodonts, brontotheres, and massive sixhorned uintatheres (Gazin, 1976; Grande and Buchheim, 1994; Gunnell and Bartels, 1994; McGrew and Casilliano, 1975; Murphey et al., 2001). Fossils and Biochronology of the Bridger Formation One of the world’s most abundant and diverse middle Eocene vertebrate faunas is preserved in the Bridger Formation. More than 86 species representing 67 genera, 30 families, and 13 orders of fossil mammals are recognized (Gazin, 1976). Joseph Leidy’s 1869 description of Omomys carteri was the first scientific description of a fossil from the Bridger Formation. Subsequently, Bridger fossils have been the subject of numerous publications, including many classic papers by pioneers of American vertebrate paleontology (Cope 1872, 1873; Granger, 1908; Leidy, 1869, 1871, 1872a; Marsh, 1871, 1886; Matthew, 1909; Osborn, 1929). Like many other highly fossiliferous formations, the Bridger contains an abundance and diversity of fossils that make it well suited for paleontological research, most of which has focused on the phylogenetics, systematic paleontology, and biostratigraphy of the vertebrate fauna (Covert et al., 1998;

Middle Eocene rock units in the Bridger and Uinta Basins Evanoff et al., 1994; Gazin, 1957, 1958, 1965, 1968, 1976; Gunnell et al., 2009; Krishtalka et al., 1987; McGrew and Sullivan, 1970; Robinson et al., 2004; West and Hutchison, 1981). Preserved in a variety of sedimentary environments, preservational states, associations, and in locally varying abundances, these fossils include primarily vertebrates and mollusks, with less common plants and ichnofossils. Plant fossils include leaves, seeds, and wood, which is sometimes algal covered (see Murphey et al., 2001). Ichnofossils include solitary bee cases, earthworm pellets, caddis fly larvae, and fish pellets. Vertebrate fossils include fish, amphibians, reptiles (lizards, snakes, turtles, and crocodilians), a diversity of birds (see Murphey et al., 2001), and mammals. Mammalian fossils include apatotheres, artiodactyls, chiropterans, carnivores, condylarths, dermopterans, dinoceratans (uintatheres), edentates, insectivores, leptictids, marsupials, pantolestids, perissodactyls, primates, rodents, taeniodonts, and tillodonts (Gazin, 1976; Woodburne et al., 2009a, 2009b; unpublished paleontological data, University of Colorado Museum, compiled in 2002). Despite the relative ease with which diverse and statistically significant fossil samples can be collected, and the large historical collections of Bridger vertebrates available in many museums, taphonomic and paleoecologic studies of Bridger vertebrate faunas are relatively few (Alexander and Burger, 2001; Brand et al., 2000; Gunnell, 1997; Gunnell and Bartels, 1994; Murphey et al., 1999; Murphey and Townsend, 2005; Townsend, 2004; Townsend et al., 2010). Over the past twenty years, stratigraphically documented fossil collections made by workers from the University of Colorado Museum, Denver Museum of Nature and Science, University of Michigan Museum of Paleontology, and more recently by the San Diego Natural History Museum, have added significantly to existing biostratigraphic knowledge of the Bridger Formation. These collections, together with precise provenance data, have made it possible to define formal biochronologic units for the Bridgerian NALMA, most of which are based upon stratotype sections that are located in the Bridger Formation. Gunnell et al. (2009) have divided the Bridgerian into four “biochrons.” Formerly referred to as Gardnerbuttean land mammal sub-age, or Br0, biochron Br1a is the only Bridgerian biochron not found in the Bridger Formation. Its stratotype section is the Eotitanops borealis interval zone of the Davis Ranch section of the Wind River Formation. Biochron Br1b is equivalent to the lower Blacksforkian, and its stratotype spans the Bridger A (lower part of the Blacks Fork Member). Biochron Br2 is equivalent to the upper Blacksforkian, and its stratotype section spans the Bridger B (upper part of the Blacks Fork Member). Biochron Br3 is equivalent to the Twinbuttean, and its stratotype section spans the entire Bridger C and D (Twin Buttes Member). The uppermost member of the Bridger Formation the Turtle Bluff Member, or Bridger E, is the stratotype section for the earliest Uintan biochron, Ui1a (Gunnell et al., 2009; Walsh and Murphey, 2007). In summary, the mammalian fauna of the Bridger Formation has been used to formally define biochrons Br1b, Br2, Br3, and Ui1a. Appendix A is a biochronologic range chart for Bridgerian, Uin-

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tan and Duchesnean mammalian taxa based on the most up to date information (modified from Gunnell et al., 2009). The fossil assemblages of the Bridger Formation and other Eocene rock units in the greater Green River basin provide an unprecedented opportunity to study ancient communities and environments. Studies of these fossils and the rocks in which they are preserved are the source of much of our knowledge of the Eocene Epoch of North America. The vertebrate faunas are of particular scientific importance because they represent an exceptional record of early Tertiary mammalian evolution and diversification spanning the Wasatchian, Bridgerian, and earliest Uintan NALMAs. Bridger Formation Field Trip Stops The field trip route travels through the Bridger basin in an approximately stratigraphic manner. After leaving historic Fort Bridger, the staging area for many of the early fossil collecting expeditions to the Bridger Formation, the route travels east along Interstate 80 crossing through badland outcrops of the Bridger B that are stratigraphically close to the Bridger A-B boundary (Blacks Fork Member). We examine the base of the Bridger B, defined by the Lyman limestone, near Little America. Traveling back west along I-80, we then visit exposures of the Bridger B near historic Church Butte. We then continue west to Lyman and then head south along Wyoming State Highway 414 to the historic Grizzly Buttes badlands in the Bridger B. We continue south along Wyoming 414, climbing stratigraphically through the Bridger C and D (Twin Buttes Member), and examine exposures of this interval in the vicinity of Sage Creek and Hickey mountains, and at the “Lonetree Divide” (base of Bridger D). Weather permitting, we will then make our way to the southwest rim of Cedar Mountain and visit exposures of the Bridger E (Turtle Bluff Member). Finally, we will head east along Highway 414 along the south side of Cedar Mountain with excellent vistas of the Bridger C, D, and E that are overlain by the Oligocene Bishop Conglomerate. The field trip concludes after visiting exposures of the Bridger C at the base of Black Mountain. Note that all field trip distances are provided in statute (miles), whereas stratigraphic thicknesses are provided in both statute and metric units. All distances were measured using a handheld GPS device calibrated to the NAD27 datum. (Note that throughout this field trip guide, the mileages given refer to the prior stop unless specified otherwise.) Stop 1. Fort Bridger State Historic Site Parking Lot (0.0 mi; cumulative 0.0 mi) Fort Bridger was originally established as a trading post in 1843 by trapper Jim Bridger and his guide Louis Vasquez. The U.S. Army acquired the trading post in 1857 during the Mormon War. It was located along the emigrant trail to Oregon, California, and Utah, and more than twenty years after the establishment of the trading post, the route of the newly constructed Union Pacific Railroad passed not far to the north. As discussed

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in greater detail above, many of the early scientific expeditions to the Bridger Formation were based out of Fort Bridger. Yale University paleontologist O.C. Marsh and his field classes stayed at Fort Bridger before heading out to the Bridger badlands in 1870, 1871, and 1873. Rival paleontologist E.D. Cope stayed at the fort in 1872 during his only fossil collecting expedition to the Bridger. Joseph Leidy, often regarded as the father of North American Vertebrate Paleontology and the first paleontologist to formally describe a Bridger Formation fossil, made his only fossil collecting trip to the west in 1872, and also stayed at Fort Bridger. Today, Fort Bridger is a state historical site and has been partially reconstructed. The low butte just to the west of Fort Bridger is Bridger Butte, which is capped with Quaternary gravels and is composed of Bridger B strata. Turn west out of the fort parking lot along the Interstate 80 business loop and then turn east onto I-80 (toward Green River). Drive for ~32 mi and take the Granger Junction exit (Exit 66) heading north along U.S. Highway 30. Follow U.S. 30 for 1.8 mi after exiting the interstate. Then turn east and cross the cattle guard onto a dirt road for 0.2 mi at which point you will arrive at the route of old Highway 30 (unmarked gravel road that is still paved in places). Park immediately after turning right (southeast) onto old Highway 30. Outcrops of the Lyman limestone are located just to the east. Stop 2. The Lyman Limestone at Granger Junction (36.3 mi from Stop 1; cumulative 36.3 mi) This stop provides a close-up look at the Lyman limestone, which marks the boundary between the Bridger A and the lower Bridger B within the Blacks Fork Member (Figs. 3 and 4). Here, the Lyman limestone is a gray limestone with locally abundant shells of the gastropod Goniobasis. The presence of this highspired snail is a useful diagnostic indicator for this marker unit at many localities in the Bridger basin. The Lyman limestone is widespread in its distribution. It is exposed to the west where it forms the bench that is visible to the south of I-80 upon which its namesake, the town of Lyman, is situated, almost as far east as the Rock Springs uplift, at least 15 mi north of Granger, and almost as far south as the town of Manila, Utah. Stratigraphically below the Lyman limestone are strata of the Bridger A. This interval has been problematic for paleontologists because it is sparsely fossiliferous. P.O. McGrew and R. Sullivan worked on the stratigraphy and paleontology of the Bridger A in the late 1960s and published the results of their work in 1970. More recently, Gregg Gunnell and colleagues from the University of Michigan Museum of Paleontology have greatly expanded the known diversity of the Bridger A (Gunnell, 1998; Gunnell and Bartels, 2001). This has made possible the recent formalization of new biochronologic units (Gunnell et al., 2009). As discussed above in “Fossils and Biochrononlogy,” the Bridger A contains a mammalian fauna (biochron Br1b) that is biostratigraphically distinct from the fauna of the Bridger B (Br2).

Optional Stop Approximately 18 m (59 ft) stratigraphically above the Lyman limestone 1.7 mi to the southeast along Old Highway 30 is an unusual type of deposit for the Bridger Formation. Park approximately one-third of the way up the hill and look for abundant dark-brown rock fragments littering the slopes underlain by a thick green mudstone interval. The thin, dark-brown bed contains abundant fossil caddis-fly larval cases and other more enigmatic fossils preserved in what appear to be algal covered logs (SDSNH Loc. 5783). The taphonomy and paleoecology of this unit has yet to be adequately studied. The fossil-bearing bed is overlain by a 2.8 m (9 ft) thick sequence of green to tan, well-indurated, platy, fine-grained, silty sandstone. It is underlain by a 1.5 m (5 ft) thick, platy, grayish-brown, non-fossiliferous, mudstone with a distinct top contact. Insect and plant fossils are sparse in the Bridger Formation, and this bed contains the most abundant insect fossils known from the formation. Return to I-80 and head west. Heading west along I-80, the first prominent butte you come to south of the interstate is Jagged Butte, which is capped by the Jagged Butte limestone. The second prominent butte you come to (approximate highway milepost 56.5) is Wildcat Butte, which is capped by the Sage Creek limestone (Sage Creek white layer of Matthew, 1909), and which forms the base of the Twin Buttes Member. Exit I-80 at the Church Butte exit (Exit 53) and turn north onto Church Butte Road (no sign). At 19.2 mi from Stop 2, with Church Butte just to the east of your location, turn left (southwest) onto Granger Road, Uinta County Road (CR) 233. At mile 21.0, turn southeast off of CR 233 onto a two-track road heading toward the westernmost point of Jackson Ridge. Park where the two-track road crosses the pipeline right-of-way at 21.2 mi from Stop 2. Stop 3. Church Butte and Jackson Ridge (21.2 mi from Stop 2; cumulative 57.5 mi) Church Butte is a large-linear badland knob formed by the erosion of rocks of the middle Bridger Formation (lower Bridger B beds; Figs. 4, 5 and 7A). The butte was a landmark along the old Oregon-California-Mormon Trail, now Uinta County Road 233. Just to the west of the butte is a north-south–trending rim separating Porter Hollow on the east with the valley of the Blacks Fork River on the west. The trail dropped off the rim just to the southwest of Church Butte, and the outcrops of Bridger Formation below the rim were easily accessed by early geologists and paleontologists who traveled along the trail. Church Butte and the rim exposures are all within the middle part of the Blacks Fork Member of the Bridger Formation of Wood (1934), or the lower Bridger B of Matthew (1909). The rocks in the area are primarily interbedded brown to green mudstone sheets and brown to gray sandstone ribbons and sheets. The sequence includes two sandstone-dominated intervals and three mudstone-dominated intervals (Fig. 6). Four thin but regionally widespread marker units occur in the sequence that is 120 m (394 ft) thick. The following descriptions are of the Bridger exposures along the west side of the rim, over an area

Middle Eocene rock units in the Bridger and Uinta Basins approximately three square miles south of where the county road crosses the rim. The two sandstone-dominated intervals are characterized by a series of thick ribbons to broadly lenticular sheet sandstone bodies within a sequence of stacked, thin, muddy sandstone and mudstone sheets. Sandstones can comprise 100% of the total thickness within the sandstone-dominated interval, but the lower interval averages 58% sandstone and the upper interval averages

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84% sandstone within a minor amount of mudstone. The thick sandbodies within these sandstone-dominated sequences are highly connected both laterally and vertically. The thick sandbodies have a reticulate pattern, with some sandbodies oriented toward the south-southeast (vector mean of 173°) and others oriented toward the east-southeast (vector mean 118°). The sinuosities of the individual sandbodies are low (mean 1.02). These sandstone-dominated intervals represent a river system with

Figure 4. Generalized stratigraphic section of the Bridger Formation in the southern Green River basin, southwestern Wyoming. Isotopic ages reported by Murphey et al. (1999) have been recalculated using the current 28.201 Ma sanidine standard for the Fish Canyon Tuff (Renne et al., 1998).

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Figure 5. Index map of the western Bridger basin, Uinta and Sweetwater counties, Wyoming.

Middle Eocene rock units in the Bridger and Uinta Basins

Figure 6. Geologic map of the Church Butte–Jackson Ridge area showing major marker beds in the lower Bridger B interval, laterally extensive sandstone sheet intervals, and trends of major sandstone channel-belt deposits. Also shown are important sites of the Hayden 1870 expedition, including known sites where W.H. Jackson took photos. See the text for details.

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numerous splays and local avulsions. The two intervals outcrop as cliffs in the badland exposures. The mudstone-dominated intervals are characterized by thick ribbon sandstone bodies that are typically separated from adjacent sandstones by extensive mudstone beds. Mudstonedominated intervals have sandstone contents that range from 10% to 35% of total interval thickness. The sandstone ribbons represent large channels carrying mostly medium sand within a mud-dominated system. The paleocurrent indicators in these ribbons (mostly medium- to thick-trough crossbed sets) and sandbody orientations indicate an original flow toward the eastsoutheast (vector mean of 120°). The sinuosities of the sandstone bodies are low (mean 1.03) and their geometry is in a “broken stick” pattern with long straight reaches and short sharp bends. Sandbody widths and thicknesses are relatively small in straight reaches, but at bends the sandbodies are thicker and wider and contain well-developed lateral accretion sets. Fossil bones typically accumulate near the bases of these bends. Thin sandstone sheets representing overbank splay deposits are rare and are limited to near their source channels. The mudstonedominated intervals outcrop as benches and slopes in the badland exposures. The Eocene streams which deposited the lower Bridger B channel sandstones in this area were perennial and flooded every year. This is indicated by the abundance of freshwater turtles, gar-pike scales (and other fish bones), and a large freshwater snail fauna in the overbank deposits. The channel-belt deposits also contain the shells of numerous freshwater mussels (unionid clams), which indicate perennial, well-oxygenated waters in streams and rivers. Fossil plants of this time (MacGinitie and Leopold, 1972) indicate subtropical temperatures and mesic moisture with seasonal precipitation. There are four regionally widespread marker beds in the Bridger exposures in this area. Two widespread thin limestone sheets occur at the base and top of the section in the Church Butte area. The lower limestone is the Lyman limestone at the base of the Bridger B (along the Blacks Fork), and in this area it is a brown to gray ostracodal limestone with scattered catfish bones. The upper limestone occurs on the flat-surface on top of the rim, just south of the county road. This upper limestone is a brown micrite with brown to black banded chert masses and scattered large planorbid snail shells (Biomphalaria sp.). Both limestone beds can be mapped over much of the Bridger basin in lower Bridger B exposures. The predominantly fluvial sequence preserved in the Church Butte area was bracketed by these widespread lacustrine deposits. Two lithified volcanic ashes (tuffs) occur in the section. The lower tuff is represented by a red clayey mudstone that ranges from 0.2 to 0.6 m (0.6 to 2 ft) thick, 33 m (108 ft) above the Lyman limestone. The bed does not contain euhedral crystals in this area, but in other parts of the Bridger basin this bed thickens and is white with euhedral biotite crystals. This bed has not been radiometrically dated. This red tuff has been mapped over much of the western Bridger basin. A second tuff bed occurs between

10.1 and 11.7 m (33 and 39 ft) below the upper limestone and ranges in thickness from 0.5 to 0.7 m (1.6 to 2.3 ft) thick. It is a tan to olive clayey mudstone bed that weathers gray and contains abundant euhedral crystals of biotite and hornblende. Sanidine in this tuff has produced a 40Ar/39Ar age of 48.27 Ma (Murphey et al., 1999, given as 47.96 ± 0.13 Ma) recalculated using the current 28.201 Ma sanidine standard for the Fish Canyon Tuff (Renne et al., 1998). This upper tuff is called the Church Butte tuff, with the type locality located at the north side of the point on the east end of the long ridge called Jackson Ridge (UTM coordinates of Zone 12T, 572123mE, 4592105mN, WGS 84). The Church Butte tuff occurs throughout the Bridger basin wherever lower Bridger B rocks are exposed. Many of the first fossils collected from the Bridger Formation came from the Church Butte area. The first geologist known to have collected fossils from the area was Ferdinand V. Hayden. In 1868 he collected fragments of a fossil turtle that were later described as Trionyx guttatus by Leidy (1869). Hayden returned to the area as part of the Geological and Geographical Survey of the Territories of 1870. The survey camped just to the west of the area along the Blacks Fork River and collected fossils in the Church Butte area on September 10 and 11 of 1870 (Hayden, 1872, p. 41). The 1870 survey was the first time the pioneer photographer W.H. Jackson accompanied Hayden. Years later, Jackson recalled these two days: Twelve miles farther on we came to Church Buttes, a remarkable formation in the Bad Lands and a famous landmark along the old trail. While Gifford [an artist of the 1870 expedition who assisted Jackson] and I were making pictures of the interesting scenes, the geologists under the lead of Dr. Hayden were digging for fossils. They collected a wagon load of ancient turtles, shell fish, and other creatures…. For my part, I made seventeen negatives during the day, something of a record for wet plate work, considering the many changes of location I had to make in getting the different views. (Jackson and Driggs, 1929, p. 89, 91)

The best known of Jackson’s photos from the area (Fig. 7B) was taken near the end of a long badlands ridge that is herein named Jackson Ridge in honor of the photographer. The upper limestone bed marker in the area is named the Jackson Ridge limestone. Jackson’s photos of the area document the type area of such fossil mammals as Notharctus tenebrosus, Palaeosyops paludosus, Hyrachyus agrestis, and Microsus cuspidatus, all described by Leidy (1870, 1872b) and illustrated in 1873. The mollusk type specimens collected at Church Butte by the Hayden survey include Physa bridgerensis, “Viviparus” wyomingensis (a land snail that is similar in form to the aquatic Viviparus), and “Unio” leanus described by Meek (1870, 1871, 1872). Seven other species of fossil mammals have their type area in or near the Church Butte area, and these were collected by such paleontologists as E.D. Cope, O.C. Marsh, and J. Wortman. Type species of fossil mammals collected from the Church Butte area are listed in Table 1.

Middle Eocene rock units in the Bridger and Uinta Basins Drive back onto CR 233, and turn left (southwest) toward Lyman. Turn left at mile 5.5 onto CR 237 which then crosses the Blacks Fork River, winding south to I-80. Turn westbound (toward Evanston) onto I-80 at mile 7.4. Pass the Lyman exit and drive to the Mountain View-Fort Bridger (Exit 39, 15.5 mi from Stop 3). Turn south onto Wyoming State Highway (SH) 414, crossing the Blacks Fork River and climbing up onto the Lyman limestone at the top of the hill. Continue through Urie and Mountain View, where the highway will bend to the east near the center of town. Refer to Figure 8 for a map that shows the major geographic features of the remainder of the field trip route. As you drive east from the center of Mountain View along Highway 414, the badlands to the south that are visible beginning at SH 414 milepost 105 were known to the early residents and explorers as “Grizzly Buttes” (lower and middle Bridger B). The north end of the badlands to the northeast constitute the type area of

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the Blacks Fork Member. Continue southeast on Highway 414 and the highway rises onto the Cottonwood Bench. Immediately after reaching the top of this bench, at 29.7 mi from Stop 3, turn east and then immediately north. At 0.2 mi from the turn off, do not turn east on Burnt Fork Road (Bureau of Land Management Road [BLM] 4315) and instead continue traveling north. At 30.4 mi from Stop 3, turn west onto the two-track road and follow it for 0.6 mi to the Grizzly Buttes overlook. Stop 4. Grizzly Buttes (31.0 mi from Stop 3; cumulative 88.5 mi) Heading southeast from Mountain View, Wyoming State Highway 414 rises through a panel of badland exposures and climbs onto a high flat, called the Cottonwood Bench. The bench is capped by gravels derived from the Bishop Conglomerate and transported to the area by Cottonwood and Sage Creeks. Below the gravel-flat is a series of badlands cut by Leavitt Creek, Little

Figure 7. (A) William H. Jackson photo of Church Butte taken on 10 or 11 September 1870. View is to the east northeast. UTM location of the photo site is Zone 12T, 571955mE, 4595101mN, WGS84 datum (U.S. Geological Survey photo jwh00462). (B) William H. Jackson photo of the west end of Jackson Ridge, taken mid-day either on 10 or 11September 1870. The view is toward the northwest, and includes the Hayden Survey campsite along the Blacks Fork River. Notice the crack that was in the original glass-plate negative. The UTM location of the photo site is Zone 12T, 571318mE, 4592360mN, WGS84 datum (U.S. Geological Survey photo jwh00309).

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Murphey et al. TABLE 1. TYPE SPECIES OF FOSSILS FROM THE CHURCH BUTTE–BLACKS FORK AREA MAMMALS Marsupialia Peratherium innominatus (Simpson) 1928 Pantolesta Pantolestes longicaudus Cope 1872 Primates Notharctus tenebrosus Leidy 1870 Tillodontia Tillodon fodiens (Marsh) 1875 Rodentia Microparamys minutus (R.W. Wilson) 1937 Sciuravus bridgeri R.W. Wilson 1937 Carnivora Miacis parvivorus Cope 1872 Hyaenodontida Sinopa major Wortman, 1902 Condylarthra Hyopsodus paulus Leidy 1872 Perissodactyla Orohippus major Marsh 1874 Palaeosyops paludosus Leidy 1870 Cetartiodactyla Microsus cuspidatus Leidy 1870 TURTLE Trionychidae “Aspideretes” guttatus Leidy 1869 MOLLUSKS Bivalvia, Unionidae “Unio” leanus Meek 1870 Gastropoda, Pulmonata Physa bridgerensis Meek 1872 “Viviparus” wyominensis Meek 1871 Note: Data compiled from Leidy (1872a); Meek (1872); Henderson (1935); and Gazin (1976).

Dry Creek and their tributaries. The badland hills directly west of the overlook comprise the traditional “Grizzly Buttes” of the early explorers, but the name is not known to the modern population of the Smith’s Fork valley (see history of paleontological investigations). Matthew (1909, p. 297) stated about the buttes: “This is the richest collecting ground in the basin; thousands of specimens have been taken from it, and many skulls and skeletons more or less complete.” Type species of fossil mammals collected from the Grizzly Buttes area are listed in Table 2. The lower half of the Bridger B is exposed in the Grizzly Buttes and along the Cottonwood Bench escarpment. Not far below the Quaternary gravels at this overlook is a widespread limestone that was named by Matthew (1909) the Cottonwood white layer (now known as the Cottonwood limestone). It is a white micritic limestone that is very widespread but is locally absent in the Church Butte area. The Cottonwood limestone is typically 5 m (16 ft) above the Church Butte tuff, but in this area it is 10.4 m (34 ft) above the tuff. The thickness of intervals between the widespread marker beds increases from the Church Butte area toward the southwest. The Jackson Ridge limestone has been eroded by Cottonwood Creek on the bench, but in this area it is typically 6 m (20 ft) above the Cottonwood White Layer.

The Church Butte tuff is a prominent gray band about half way down the escarpment. Notice that channel sandstones are not as abundant in the lower Bridger B rocks below you as they are in the Church Butte area. To the east is a prominent escarpment rising far above the Cottonwood Bench. This escarpment is capped by the Sage Creek White Layer, the boundary between the Blacks Fork and Twin Buttes members of the Bridger Formation (the boundary between Matthew’s Bridger B and C). Almost all the upper half of the Bridger B is exposed in the west face of the escarpment. Return to Wyoming State Highway 414 and travel north for 5.7 mi. Then turn east and drive for 0.2 mi and park on the north side of the road. A short walk to the northeast will lead you to Sage Creek limestone and the type locality of the Sage Creek white layer. Stop 5. Sage Creek White Layer Type Locality (5.9 mi from Stop 4; cumulative 94.4 mi) This outcrop of the Sage Creek white layer is located next to site of the old Sage Creek stage station and Sage Creek Spring along the old Lonetree stage road. It was first described and photographed by Sinclair in 1906 (Fig. 9), and then named and mapped by Matthew (1909). The Sage Creek white layer is the base of Matthew’s Bridger C, the base of the Twin Buttes Member, and the base of the upper Bridger Formation as presently defined. Since Matthew’s (1909) work, this unit has been renamed the Sage Creek limestone, and is the base of the lower Bridger C of Evanoff et al. (1998), Murphey (2001), and Murphey and Evanoff (2007). The general stratigraphy of the upper Bridger Formation in the Sage Creek Mountain area is illustrated in Figure 10. At its type locality, the Sage Creek limestone is 4.1 m (13.5 ft) thick. It consists of a lower massive tan micritic limestone, a middle shaly limestone with dark-gray to black chert bands, and an upper platy to shaly limestone. Elsewhere, it includes massive to blocky marly and micritic limestone, ledgy marlstone, and platy calcareous shale, and is locally interbedded with green to brown mudstone and claystone and thin carbonaceous shale. Fossils of this unit consist of scattered gastropods, bone fragments (mostly fish), and turtle shell fragments, and the limestone within it is locally stromatolitic. The Sage Creek limestone supports a very widespread bench, and it is the thickest and most widespread lacustrine deposit in the upper Bridger Formation. Stratigraphically overlying the Sage Creek limestone within the lower Bridger C are two other limestone beds that are much thinner but are also widespread: the Whisky Reservoir limestone and the Butcher Knife limestone (see Fig. 10). The lower Bridger C is the least fossiliferous subunit of the upper Bridger Formation (Twin Buttes and Turtle Bluff members), despite the fact that it is by far the most geographically widespread. Continue south along Highway 414 for 3.2 mi. Traveling south, the highway route travels up section through the lower Bridger C and into the middle Bridger C. Sage Creek Mountain is the highest point on the west side of the highway and Hickey Mountain is the highest point on the east side of the highway.

Middle Eocene rock units in the Bridger and Uinta Basins Both of these mountains are capped by the Oligocene Bishop Conglomerate. At 3.2 mi from Stop 5, pull into the Henry #1 gas well pad on the east side of the road. Stop 6. Soap Holes and Hickey Mountain and Limestones (3.2 mi from Stop 5; cumulative 97.6 mi) The Soap Holes limestone, the lower of the two thin rusty brown limestone beds visible at this cliffy exposure, is a widespread marker unit that forms the base of the middle Bridger C (Evanoff et al., 1998; Murphey, 2001; Murphey and Evanoff, 2007). It is believed that Matthew (1909) considered this bed to be equivalent to his Burnt Fork limestone, which is a lithologically similar unit that is exposed to the southeast in the Henrys Fork Valley but is actually not present in the section in this part of the basin. In the Henrys Fork Valley, however, it is in fact is 33 m (108 ft) higher than the Soap Holes limestone. The Soap Holes limestone contains few fossils, but it is noteworthy that it is stratigraphically closely associated with fossil logs at several localities. Fossils of the Soap Holes limestone include isolated, disarticulated and poorly preserved bones of fish, reptiles (especially turtles), and mammals within and on top of the unit. In the Black Mountain area it is locally underlain by thin carbonaceous shale beds which preserve plant fragments. The Sage Creek and Soap Holes limestones have in fact yielded the fewest vertebrate fossils of any upper Bridger lacustrine deposits. Situated within the middle Bridger C 10.5 m (34 ft) above the base of the Soap Holes limestone (in the upper Bridger

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Formation reference section, Murphey and Evanoff, 2007), the Hickey Mountain limestone is a well studied and very important unit paleontologically. It has a relatively limited areal distribution, occurring over a distance of ~5.6 mi north of Hickey Mountain and west of Sage Creek Mountain, and is the upper limestone bed exposed on the cliff at this stop. This unit provides an excellent example of one of the most paleontologically prolific depositional settings in the upper Bridger Formation. The early fossil collectors were the first to notice the close association between vertebrate fossils and the “white layers,” which are typically limestone and marlstone beds that were deposited in shallow lakes and ponds. More recently, paleontologists observed that it is not the marlstone beds that contain the majority of vertebrate fossils, but the immediately overlying and underlying mudstone beds. These mudstones, which are occasionally carbonaceous, are inferred to have been deposited along lake margins during lake transgressions and regressions (Murphey, 1995; Murphey et al., 2001). Typically, the limestone and marlstone beds contain the remains of mostly aquatic organisms such as snails, clams, fish, amphibians, pond turtles, and crocodilians. The lake margin mudstones contain a mixed aquatic and terrestrial assemblage, and the terrestrial elements include locally abundant reptiles such as lizards, as well as bird bones and mammal bones and teeth. One particularly prolific fossil locality, the Omomys Quarry, is located approximately ½ mi west of this stop in the Hickey Mountain limestone and overlying

Figure 8. Map of a part of the southern Green River basin encompassing most of the area of outcrop of the upper Bridger Formation. Map shows both modern and historic geographic terminology (from Murphey and Evanoff, 2007).

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mudstone. This unusual fossil accumulation has produced over 2,300 specimens of vertebrates, gastropods, and plants from an 8–10-cm-thick deposit in a 4 m2 area (Murphey et al., 2001). What makes the assemblage so unusual is that it contains a high concentration of dental and post-cranial remains of the primate Omomys, avian skeletal remains, and eggshell fragments. The unusual components of the fauna are superimposed on a more typical Bridger fauna that occurs at the quarry and lateral to it in the same stratigraphic interval. Four taphonomic agents have been postulated for the formation of the Omomys Quarry fossil accumulation: (1) an attritional accumulation of aquatic taxa in lacustrine sediments; (2) an attritional accumulation of both aquatic and terrestrial taxa in shoreline sediments; (3) an attritional accumulation consisting primarily of bird bones and eggshell formed in close proximity to a nesting area; and

TABLE 2. TYPE SPECIES OF FOSSIL MAMMALS FROM GRIZZLY BUTTES Lipotyphla Entomolestes grangeri Matthew 1909 Nyctitherium serotinum (Marsh) 1872 Nyctitherium dasypelix (Matthew) 1909 Plesiadapiformes Mycrosyops elegans (Marsh) 1871 Primates Smilodectes gracilis (Marsh) 1871 Tillodontia Trogosus castoridens Leidy 1871 Pholidota Metacheiromys marshi Wortman 1903 Metacheiromys tatusia Osborn 1904 Metacheiromys dasypus Osborn 1904 Rodentia Thisbemys plicatus A.E. Wood 1962 Leptotomus parvus A.E. Wood 1959 Reithroparamys delicatissimus (Leidy) 1871 Pseudotomus robustus (Marsh) 1872 Ischyrotomus horribilis A.E. Wood 1962 Mysops minimus Leidy 1871 Mysops parvus (Marsh) 1872 Sciuravus nitidus Marsh 1871 Tillomys? parvidens (Marsh) 1872 Hyaenodontida Sinopa rapax Leidy 1871 Sinopa minor Wortman 1902 Tritemnodon agilis (Marsh) 1872 Limnocyon verus Marsh 1872 Carnivora Thinocyon velox Marsh 1872 Viverravus gracilis Marsh 1872 Oödectes proximus Mattehw 1909 Vulpavus profectus Matthew 1909 Perissodactyla Palaeosyops major Leidy 1871 Limnohyops priscus Osborn 1908 Helaletes nanus (Marsh) 1871 Cetartiodactyla Helohyus plicodon Marsh 1872 Note: Compiled from Gazin (1976).

(4) a predator accumulation dominated by Omomys but probably including other vertebrates formed by owls in close proximity to a nest, day roost or night feeding station. The fauna and flora of the Omomys Quarry is listed in Table 3. The same pattern of fossil distribution observed in the Hickey Mountain limestone occurs throughout the upper Bridger Formation as illustrated in Figure 11. Most fossils are found in association with lacustrine deposits, although stream channels are also productive. Least productive are the volcaniclastic mudstone and claystone beds that were deposited on low relief floodplains, and, together with stream channel deposits, comprise 95% of the total thickness of the upper Bridger. Examples of the floodplain deposits, here consisting of green and gray mudstone and claystone beds, are well exposed at this stop above and below the Soap Holes and Hickey Mountain limestones. Both the Soap Holes and Hickey Mountain limestones are better exposed, with some minor faulting, on the east side of the highway just to the north of this location. Continue south on Highway 414 for 1.1 mi and turn east onto the gas well road. Follow this road to the east and it will bend to the north for a total distance of 1.6 mi from Stop 6. Park on the north side of the Henry #10 well pad. The Henrys Fork limestone (type locality of this unit) and the underlying Henrys Fork tuff are exposed above the well pad on the badland hill just to your north. Look for the gray weathered bed near the top of the badland slope and an overlying thin light-gray marlstone, and bring a shovel to examine the tuff. Stop 7. Type Locality of the Henrys Fork Limestone (1.6 mi from Stop 6; cumulative 99.2 mi) The Henrys Fork limestone (and associated shore margin deposits) is another highly fossiliferous unit, and has produced hundreds of fossil mollusks and vertebrates across its distribution. It is quite widespread, covering an area of ~402 km2 (250 mi2), and was deposited in an elongate east-west–trending basin which formed in the downwarp along the Uinta Mountain front. At this location, which is near the western edge of ancient Henrys Fork Lake, the deposit is only 3 cm (1.2 in) thick, but it attains a maximum thickness of 1.65 m (5.4 ft) on the south side of Cedar Mountain near the center of its depositional basin. It is of taphonomic interest that the upper Bridger Formation with its abundant vertebrate fossils preserved in lacustrine and associated shore margin deposits contains few articulated skeletons or even partially articulated vertebrate remains, most of which have been collected in the Bridger B (see Alexander and Burger, 2001). Immediately underlying the Henrys Fork limestone is the Henrys Fork tuff, a unit that was first discovered by Emmett Evanoff while conducting field work in the Sage Creek Mountain area in 1991. This ash-fall tuff is the most analyzed tuff in the Bridger Formation, and it is beyond the scope of this paper to report the various ages that have been published. However, Murphey et al. (1999) and Murphey and Evanoff (2007) reported a 40Ar/39Ar age of 46.92 ± 0.17 Ma based on single

Middle Eocene rock units in the Bridger and Uinta Basins

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Figure 9. Photograph of the Sage Creek white layer taken by W.H. Sinclair in 1906 (Sinclair, 1906, plate 38). Note the unit numbers penned in near the left edge of the photo.

Figure 10. Generalized stratigraphic section of the upper Bridger Formation in the Sage Creek Mountain area, Uinta County, Wyoming. The diagram shows widespread and more localized markers, as well as informal submembers of Matthew (1909). Thicknesses taken from the reference section of the Twin Buttes Member and the type section of the Turtle Bluff Member (from Murphey and Evanoff, 2007).

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TABLE 3. FAUNA AND FLORA OF THE OMOMYS QUARRY TAXA NISP * PLANTAE Division Chlorophyta Chara sp. Division Tracheophyta Dennstaedtiopsis aerenchymata (fern) 2 types of dicotyledenous wood ANIMALIA Phylum Mollusca Class Gastropoda Order Lymnophila (freshwater pulmonates) Biomphalaria sp. Omalodiscus sp. Stagnicola? spp. Physa spp. Order Geophilia (land pulmonates) Gastrocopta spp. Oreoconus? sp. Phylum Chordat a Coprolites 4 Class Osteichthyes Osteichthyes undet. 200 Orde r Amiiforme s Amia sp. 8 Order Lepisosteiformes Lepisosteus sp. 98 Order Siluriformes 14 Class Amphibia Amphibia undet. 4 Order Anura Anura undet. 4 Class Reptilia Order Chelonia Chelonia undet. 5 Trionychidae undet. 1 Echmatemys sp. 1 Palaeotheca sp. 2 Order Squamata Lacertilia undet. 117 Iguanidae undet. 3 Tinosaurus sp. 1 Saniwa sp. 1 Serpentes undet. 2 Order Crocodilia Allognathosuchus sp. 34 4 Crocodilia undet. Class Aves Order Ciconiiformes Juncitarsus gracillimus 4 Order Falconiformes Accipitridae (2 species) 7 Order Charadriiformes Burhinidae 5 Order Gruiformes Rallidae 5 Geranoididae 8 (191) Total Aves (includes undet. specimens not listed above) (continued)

TABLE 3. (continued) TAXA

NISP*

Class Mammalia Mammalia undet. 150 Order Marsupialia Peratherium sp. 13 Peradectes sp. 1 Order Rodentia Rodentia undet. 13 Paramys sp. 1 Thisbemys sp. 1 Microparamys sp. 1 Ischyromyidae undet. 1 Sciuravus sp. 9 2 Pauromys sp. Sciuravidae undet. 6 Order Apatemyida Apatemys sp. 1 Order Lipotyphla Lipotyphlan undet. 28 Scenopagus sp. 2 Entomolestes sp. 7 Centetodon sp. 6 A p t e r n o d o n ti d a e u n d e t . 1 Nyctitherium sp. 4 Order Plesiadapiformes Uintasorex sp. 1 Order Primates Notharctus sp. 2 Omomys sp. nov. 214 Order Condylarthra 22 Hyopsodus sp. Order Cetartiodactyla Cetartiodactyla undet. 1 Homacodon sp. 1 Order Perissodactyla Hyrachyus sp. 1 TOTAL 1,183 Note: Eggshell is not included. From Murphey et al. (2001). *Number of identifiable specimens for vertebrates.

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crystal laser fusion analysis of sanidine, plagioclase, and biotite. Recalculated using the current 28.201 Ma sanidine standard for the Fish Canyon Tuff (Renne et al., 1998), the age of the Henrys Fork tuff is 47.22 Ma. Ash-fall tuff deposits comprise less than 1% of the total thickness of the upper Bridger Formation, and based on their mineralogy, are believed not to have originated in the Absaroka volcanic field to the north like other volcaniclastic Bridger Formation sediments, but rather in the Challis volcanic field located in central Idaho. At its type locality on the south side of Cedar Mountain, the tuff is 0.95 m (3.1 ft) thick (Murphey and Evanoff, 2007). The tuff is a blocky, non-calcareous, gray to white biotitic claystone with a distinct bottom contact and a diffuse top contact. It contains biotite, zircon, allanite, and apatite crystals. Plagioclase is the most abundant feldspar. It typically consists of a structureless lower unweathered portion with coarse euhedral biotite (up to 1.3 mm in diameter) which grades upward into a reworked portion with less coarse and less abundant biotite.

The base of the Henrys Fork tuff forms the base of the upper Bridger C, and is 121 m (397 ft) above the base of the Sage Creek limestone in upper Bridger Formation reference section (Murphey and Evanoff, 2007). Weathered badland exposures of the Henrys Fork tuff form a distinctive dark-gray weathering bed that is readily discernable from other Bridger lithologies, especially when wet. Return to Highway 414 and drive south for 3.0 mi (note the exposures of Henrys Fork tuff which is visible as a subtle gray bed on both sides of the highway just above road level after turning back onto the highway). Then turn west onto the access road for Conoco Fed #20-2 gas well pad. Proceed to the well pad and park (3.2 mi from Stop 7). The route you just drove continued up section through the upper Bridger C to the level of the Lonetree limestone (base of lower Bridger D) which is at the approximate level of the highway at the Lonetree Divide. This is the stratigraphically highest point that Highway 414 attains in the Bridger Formation.

Figure 11. Stratigraphic distribution of catalogued mammalian specimens from the upper Bridger Formation (University of Colorado Museum collections). BELS—Basal Bridger E limestone; HFLS—Henrys Fork limestone; HMLS—Hickey Mountain limestone; HRLS— Hickey Reservoir limestone; LTLS—Lonetree limestone; SCLS— Sage Creek limestone; SHLS—Soap Holes limestone; ULS—Upper White limestone.

Stop 8. The Lonetree Divide (3.2 mi from Stop 7; cumulative 102.4 mi) This area provides some excellent vistas of the upper Bridger Formation and its marker units, especially from the top of the ridge just to your north. The base of the lower Bridger D, the Lonetree limestone (Lonetree white layer of Matthew, 1909), is well exposed at the base of the badland slopes at road level. The base of the middle Bridger D, the Basal blue sheet sandstone, is exposed on the slopes of the prominent conical butte to your west as well as on parts of the ridges to your north and south. The prominent butte, called “Old Hat Mountain” by the locals, is an erosional remnant of Hickey Mountain (Fig. 12A) to which it is still attached. The “rim of the hat” is the Upper White limestone (Upper white layer of Matthew, 1909). The butte is capped by a thin interval of red mudstone of the Turtle Bluff Member (Bridger E of Matthew, 1909). To your northeast is Sage Creek Mountain, with a thick sequence of Bridger E (red beds overlying gray beds of Bridger D) visible near its summit. The Basal Bridger E tuff (40Ar/39Ar age of 46.16 ± 0.44 Ma; Murphey and Evanoff, 2007) occurs just below the base of the Bridger E on Sage Creek Mountain. O.C. Marsh called Sage Creek Mountain “Big Bone Butte” because of the abundance of uintathere bones found in the area. Visible to your east is Cedar Mountain, with the thickest and best exposed sequence of Turtle Bluff Member. All three of the mountains in this area (Hickey, Sage Creek, and Cedar) are capped by Oligocene Bishop Conglomerate. Numerous fossil localities have been documented in the Lonetree Divide area. These include the classic Lonetree localities of Matthew (AMNH expeditions of 1903–1906) and Gazin (United States National Museum [now the National Museum of Natural History] expeditions between 1941 and 1969. This area was also worked by Robert M. West of the Milwaukee Public Museum during the 1970s, and by crews from the University of Colorado Museum during the 1990s. Channel sandstones in this area indicate paleocurrent directions to the southeast.

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Turn south on Highway 414 for 1.9 mi and turn east onto Cedar Mountain Rim road (BLM Road 4314). At 2.8 mi from Stop 8, the road bends to the north and travels stratigraphically through the upper Bridger C, crossing the Lonetree limestone, and continuing up through the lower and middle Bridger D. At the junction of Cedar Mountain Rim Road and Sage Creek Mountain Road (5.2 mi from Stop 8), turn south onto a twotrack road. Drive south on the two track, keeping straight at miles 6.2 and 6.3 where other tracks diverge, until you reach the Turtle Bluff Member overlook (6.4 mi from Stop 8). Note

that if the ground is wet, it is not advisable to leave the paved highway (SH 414). Stop 9. The Turtle Bluff Member on Cedar Mountain (6.4 mi from Stop 8; cumulative 108.8 mi) Looking east from this location affords an excellent view of the upper Bridger D, the highest sub-unit in the Twin Buttes Member, overlain by the Turtle Bluff Member of the Bridger Formation (Matthew’s Bridger E). The contact between the two members is shown on Figure 12B and is defined on the basis of

Figure 12. (A) View of the Bridger D and E on “Old Hat Mountain,” a prominent butte on the southeast flank of Hickey Mountain, Uinta County, Wyoming. Photo taken looking southwest. (B) View of the Bridger D and Bridger E on the southwest flank of Cedar Mountain, Uinta County, Wyoming. Photo taken looking east. BBS—Basal blue sheet sandstone; BELS—Basal Bridger E limestone; BRGB—Behunin Reservoir Gypsum bed; Tbe—Bridger E; Tbi— Bishop Conglomerate; Tbdm—middle Bridger D; Tbdu—upper Bridger D; ULS—Upper White limestone.

Middle Eocene rock units in the Bridger and Uinta Basins a limestone that occurs at the approximate level of the lowest red bed (note that some of the strata you see are slumped). The limestone that supports the bench that you are standing on is the Upper White limestone. Consisting primarily of banded red, gray, and tan beds of gypsiferous claystone and mudstone, rocks of the Turtle Bluff Member are the least volcaniclastic in the Bridger Formation. Lthologically, the Turtle Bluff Member is somewhat distinct from the rest of the formation, being similar in appearance to the red and brown sandstone, mudstone, and claystone beds of parts of the Washakie and Uinta Formations of similar age. The Turtle Bluff member occurs only on Hickey Mountain, Sage Creek Mountain, the south end of Black Mountain, and Twin Buttes, but by far the most extensive and thickest exposures occur here on the southwest side of Cedar Mountain. The type section for the Turtle Bluff Member on Cedar Mountain is a 131.5 m (431 ft) thick sequence of reddish-brown and gray claystone beds with a high gypsum content. This gypsum is both primary and secondary. Secondary gypsum consists of selenite and satin spar crystals which are abundant on the upper slopes of Cedar Mountain. Primary gypsum occurs in thin beds, but the Turtle Bluff Member on Cedar Mountain is capped by a thick and laterally extensive gypsum bed. The mostly fine-grained reddish Turtle Bluff sediments were probably derived from the adjacent Uinta Mountains based on their color, unlike those of the Bridger A–D, which were largely derived from more distal volcanic sources. The Turtle Bluff Member contains two markers: The Basal Bridger E limestone, which marks the base of the member (base of Matthew’s Bridger E), and the Behunin Reservoir Gypsum Bed, which is the youngest and stratigraphically highest well exposed rock unit in the Bridger Formation (note that Behunin is pronounced “Buhannan” by locals). Here on southwest Cedar Mountain, the Turtle Bluff Member contains four additional unnamed limestone beds, and on Twin Buttes there are three. A 2.3 m (7.5 ft) thick laterally extensive, quartz arenite bed that lies 75 m (246 ft) above the base of the member on Cedar Mountain is the only sandstone bed. Similar sandstone beds in the Turtle Bluff Member also occur on the northwest flank of Hickey Mountain and the south flank of Sage Creek Mountain, and may be roughly stratigraphically equivalent. The Behunin Reservoir Gypsum Bed is lithologically unique for the Bridger Formation. Although other gypsum beds occur in the Turtle Bluff Member, they are much thinner. Restricted to just below the southwest rim of Cedar Mountain (below the Bishop Conglomerate), this unit consists of a 7 m (23 ft) thick sequence of gray and tan unfossiliferous bedded gypsum beds interbedded with gypsiferous mudstones and marlstones. It is visible from a great distance as a prominent white bed high on Cedar Mountain. This bed is interpreted as an evaporitic playa lacustrine deposit, and may indicate changing climatic conditions near the end of Bridger Formation deposition. Because of its sparse fossils and steep, limited exposures, the biochronologic affinity of the Turtle Bluff Member has been difficult to determine. Matthew was the first worker to comment

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on the age of the member, saying that its few mammal fossils prove sufficiently that it “belongs to the Bridger Age” (Matthew, 1909, p. 296). Osborn (1929) correlated the Bridger E with the Washakie B and Uinta B, although he cited no evidence to support this correlation. Simpson (1933), Wood et al. (1941), and Gazin (1976) also regarded the Bridger E as Uintan, although this was apparently not based on fossil evidence. Based on eight isolated rodent teeth identified as Paramys cf. P. delicatior, and “several” bone fragments identified as brontothere, West and Hutchison (1981) concluded that the Bridger E (their Cedar Mountain Member) was Bridgerian. Subsequent work during the 1990s by crews from the University of Colorado Museum (Evanoff et al., 1994; Murphey and Evanoff, 2007) and in the 2000s by crews from the San Diego Natural History Museum (Walsh and Murphey, 2007) have now documented a much more diverse faunal assemblage from multiple stratigraphic levels within the Turtle Bluff Member (Table 4). Donna’s locality (UCM Loc. 92189) is located near the base of the member and is the only locality thus far to produce specimens of the newly described species of omomyid primate Hemiacodon engardae (Murphey and Dunn, 2009). Located 105 m (344 ft) above the base of the member, Roll the Bones (SDSNH Loc. 5844) and Red Lenses (SDSNH Loc. 5844) are the stratigraphically highest localities to yield identifiable fossils in the member. Hundreds of fossils have now been collected from these and other localities mostly via screenwashing of sedimentary matrix. Non-Bridgerian taxa include Epihippus, Metanoiamys, Pareumys, Triplopus, Sespedectinae indet., and Oromerycidae gen. and sp. nov. The faunal assemblage of the Turtle Bluff Member is now considered to be earliest Uintan in age (biochron Ui1a of Gunnell et al., 2009), although efforts to obtain additional fossils from this biochronologically important interval member on Cedar Mountain and other locations within the Bridger basin are ongoing. The Bridger Formation is unconformably overlain by the Bishop Conglomerate, which is visible from this stop capping Cedar Mountain. To the east of this location, it forms massive cliffs and spectacular columns. This unit is a very coarse conglomerate composed primarily of arkosic cobbles and boulders derived from the Proterozoic Uinta Mountain Group, with locally common cobbles and boulders of Paleozoic limestone (Bradley, 1964). It is as much as 40 m (131 ft) thick. The Bishop Conglomerate is unfossiliferous, but currently believed to be Oligocene in age (K/Ar 29.50 ± 1.08 Ma, biotite) based on isotopic ages obtained from a tuff that occurs within it on the south side of the Uinta Mountains (Hansen, 1986). Return to Wyoming State Highway 414, and continue south. The highway crosses the Henrys Fork of the Green River, passes by the hamlet of Lonetree, and bends to the east. There is an excellent view of the Henrys Fork tuff and Henrys Fork limestone on the north side of the highway at SH 414 milepost 128. The Henrys Fork tuff is the prominent gray bed exposed low on the slopes of Cedar Mountain not far above road level, and the Henrys Fork limestone is a prominent white

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bed immediately overlying the tuff. Continuing east, the highway passes through the hamlet of Burntfork. At highway milepost 130.6 there is a point of historic interest on the south side of the highway. This location is near the site of the first (1825) mountain man fur trading “rendezvous” led by General William Ashley and attended by a then “green” Jim Bridger and Jedediah Smith. At the McKinnon Junction (16.6 mi from Stop 9), turn north onto Sweetwater County Highway 1. Proceed to 18.8 mi from Stop 9 and pull over on the east shoulder next to the road cut. Stop 10. The McKinnon Roadcut (18.8 mi from Stop 9; cumulative 127.6 mi) This roadcut features the thickest lacustrine sequence in the upper Bridger Formation. The Sage Creek limestone is the thick blocky limestone near the top of the cut. Underlying it are at least 30 m (98 ft) of lacustrine shale, mudstone, marlstone, and limestone, and the total thickness of the lacustrine sequence is unknown. It has been postulated that this sequence and under-

lying lacustrine strata of unknown thickness represents a final transgressive phase of Lake Gosiute (Laney Shale Member of Green River Formation) (Brand, 2007), although there is little supporting evidence. Whatever the case, this sequence, combined with evidence provided by other upper Bridger lacustrine deposits (thicknesses and areal distribution), suggests that lacustrine deposition during upper Bridger deposition was most prevalent just to the north of the Uinta Mountain front. Continue north on Sweetwater County Highway 1. You will be driving through rocks of the lower Bridger C and will descend into upper Bridger B strata at approximate Sweetwater County Highway 1 milepost 16.7 before climbing stratigraphically again lower Bridger C strata at highway milepost 15.8. At mile 11.5 from Stop 10, turn east onto a two-track road toward the north end of Black Mountain. Twin Buttes is the conical peak to the south of Black Mountain. Bear right at mile 12.1. Take the left fork at mile 13.0 (look for the BLM Wilderness Study Area sign). Park at the base of Black Mountain at mile 13.8. Note that if the ground is wet, it is advisable to stay on the paved highway.

TABLE 4. TAXONOMIC SUMMARY OF THE TURTLE BLUFF MEMBER, BRIDGER FORMATION (BIOCHRON UI1A) Designation Number Taxa Index Species 2 Hemiacodon engardae Oromerycidae gen. and sp. nov. Genus LRD

6

Epihippus Oromerycidae gen. and sp. nov. Metanoiamys

Pareumys Sespedectinae indet. Triplopus

Genus HRD

7

Entomolestes Hemiacodon Oromerycidae gen. and sp.nov. Mysops

Paramys Taxymys Uintasorex

Species LRD

3

Epihippus gracilis Metanoiamys sp.

Triplopus cubitalis

Species HRD

8

Entomolestes grangeri Paramys delicatior Pauromys perditus Pontifactor bestiola

Sciuravus nitidus Taxymys lucaris Thisbemys corrugatus Uintasorex parvulus

Range-Through

36

Apatemys sp. Nyctitherium sp. Antiacodon sp. Omomys carteri Brontotheriidae spp. Orohippus sylvaticus Centetodon bembicophagus Pantolestes longicaudus Chiroptera indet. Pantolestes natans Copedelphys innominatum Pantolestes sp. Dilophodon minusculus Pauromys sp. Peradectes chesteri Harpagolestes sp. Helohyus sp. Scenopagus priscus Herpetotherium knighti Thisbemys corrugatus Herpetotherium marsupium Tillomys senex Hyopsodus sp. Trogolemur sp. Hyrachyus eximius Uintacyon vorax Isectolophus sp. Uintaparamys bridgerensis Mesonyx obtusidens Uintaparamys caryophilus Microparamys minutus Uintatherium anceps Microsyops annectens Viverravus minutus Notharctus robustior Washakius sp. Note: From Gunnell et al. (2009). LRD—lowest range datum; HRD—highest range datum.

Middle Eocene rock units in the Bridger and Uinta Basins Stop 11. Twin Buttes and Black Mountain (13.8 mi from Stop 10; cumulative 141.4 mi) Although the classic Bridger badlands and collecting areas we have already visited are located far to the west, the Twin Buttes Member and the Twinbuttean land mammal subage was named for Twin Buttes. Because of this, Murphey and Evanoff (2007) designated their type section of the Twin Buttes Member for the upper Bridger sequence on the south side of Twin Buttes, and established their Twin Buttes Member reference section of the upper Bridger for the sequence in the Sage Creek and Hickey Mountain area. However, the reference section is thicker and contains more marker units. The major stratigraphic features of the upper Bridger in the Twin Buttes and Black Mountain area are shown in Figure 13. You are standing in front of another upper Bridger marker bed. The Horse Ranch red bed occurs only in the eastern part of the basin (east side of Twin Buttes [Mass Mountain], Black Mountain and Twin Buttes). It is an ~4 m (13 ft) thick sequence of noncalcareous brick red, greenish-gray, and light-brown claystone, blocky mudstone, and blocky fine-grained muddy sandstone (Murphey and Evanoff, 2007). It is locally fossiliferous. For many years, paleontologists were vexed by the difficulty of correlating between Twin Buttes and Cedar Mountain to the west, especially considering the classic “layer cake geology” of the Bridger with very low dips and laterally persistent marker units. This was because the stratigraphic positions of the “white

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layers” did not align as expected when using the Sage Creek limestone as a datum. This problem was finally solved by locating the mineralogically diagnostic Henrys Fork tuff not far from here on Black Mountain, and using it as a stratigraphic datum. The reason that earlier workers had difficulties establishing a correlation between the Twin Buttes and Black Mountain area with exposures to the west using the “white layers” is that the lower Bridger C thins dramatically from the west to the east as evidenced on Twin Buttes, where the thickness between the Sage Creek limestone and Soap Holes limestone is 21 m (69 ft) less than in the nearest correlative sequence to the west. You are stratigraphically located within the middle Bridger C, and the Henrys Fork tuff is located 42 m (138 ft above this level). In fact, all of the major marker units present in the Twin Buttes reference section are present in the Twin Buttes type section except for the Basal blue sheet sandstone (base of middle Bridger D). The Lonetree limestone is very well exposed in the saddle between Black Mountain and Twin Buttes, and the Upper white limestone is exposed near the top of Twin Buttes. Only 21 m (69 ft) of Turtle Bluffs Member occurs at Twin Buttes, which is capped by a thin remnant of Bishop Conglomerate. The hike from this stop to the saddle between Twin Buttes and Black Mountain is well worth the effort if you have the time. This is the end of the Bridger basin portion of the field trip. From here we will head south to Manila, Utah, via Sweetwater

Figure 13. Generalized stratigraphic section of the upper Bridger Formation in the Twin Buttes area, Sweetwater County, Wyoming. The diagram shows widespread and more localized markers, as well as informal submembers of Matthew (1909). Thicknesses taken from the type section of the Twin Buttes Member, which includes the Turtle Bluff Member on Twin Buttes (from Murphey and Evanoff, 2007).

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County Highway 1 and Wyoming State Highway 414, and then continue south over the Uinta Mountains to Vernal and the Uinta basin. PART II. UINTA BASIN FIELD TRIP This portion of the field trip will examine rocks of the Uinta and Duchesne River formations in the Uinta basin. The following sections of the field trip guide provide a summary of the early Cenozoic geologic history of the Uinta basin, as well as the history of investigations, stratigraphy, depositional and paleoenvironmental history, and fossils of the Uinta and Duchesne River formations. This is followed by a detailed field trip road log. Early Cenozoic Geologic History By the late Paleocene (Clarkforkian NALMA), large shallow lakes and ponds occupied a large portion of both the Uinta and Piceance Creek basins. During the latest Paleocene and early Eocene (late Clarkforkian and early Wasatchian NALMA), there was a return to fluvial sedimentation in both basins, which is represented by rocks of the Wasatch Formation (also called DeBeque Formation by some paleontologists). This was followed by a return to dominantly lacustrine conditions in the early Eocene with the establishment of two large freshwater lakes (one in each basin). Bradley (1931) named the tongue of the freshwater lacustrine unit which was deposited during this brief time of maximum transgression the “basal tongue” of the Green River Formation where it crops out in Indian Canyon in the western part of the Uinta basin. The two early freshwater lakes in the Uinta and Piceance Creek basins appear to have been similar to the approximately contemporaneous Lake Luman, the early freshwater stage of Lake Gosiute, which existed in the greater Green River basin in Wyoming. The sediments deposited during lacustrine maximums in both the Uinta and Piceance Creek basins almost connect over the top of the Douglas Creek arch, suggesting that during at least in part of the freshwater period the two basins were hydrologically connected (Johnson, 1985). During this period, the two basins are believed to have drained externally since the lake waters in both basins remained fresh enough to support abundant freshwater mollusks (Johnson, 1985). At the end of the early Eocene (Wasatchian-Bridgerian boundary), a major transgression associated with expansion and deepening of lake water marked the beginning of Lake Uinta as an unbroken body of water across the Douglas Creek Arch in both the Piceance Creek and Uinta basins (Johnson, 1985; Moncure and Surdam, 1980). Johnson (1984) named this the Long Point transgression. After the maximum transgression, Lake Uinta extended close to the margins of the combined Uinta and Piceance Creek structural and sedimentary basins; it covered a much larger area than that of the maximum transgression during the earlier freshwater stage of the lake. Five stages in the developmental history of the larger saline Lake Uinta have been identified based on changes in water chemistry and depositional events

(Johnson, 1985). Throughout most of these five stages, the salinity of Lake Uinta increased steadily. Ultimately, salinity increased to the point at which nahcolite and halite were precipitated. The gradual filling of Lake Uinta during the middle Eocene was due to an influx of volcaniclastics from the Absaroka Volcanic Field in Wyoming to the north and from volcanic centers farther to the west, and then later by an increase in sediment resulting from rejuvenated local Laramide uplifts. There is no direct geologic evidence that Lakes Uinta and Gosiute were ever physically connected. However, fluvially deposited volcaniclastics in Lake Uinta, demonstrably derived from the Absaroka Volcanic Field in northwestern Wyoming, would be a persuasive line of evidence. This, however, would not rule out the possibility of transport of volcaniclastic material from Lake Gosiute south to Lake Uinta via streams over the lowest drainage divides after the greater Green River basin had been filled (Surdam and Stanley, 1980). As lacustrine deposition in Lake Uinta diminished because of basin filling, fluvial sedimentation resulted in deposition of the Uinta Formation in the Piceance Creek and Uinta basins (Stokes, 1986). At this time, fluvial deposition began to dominate in the basin as streams flowed into it from the north and east (Stokes, 1986). The east-west flow resulted in the formation of limey siltstone beds and fine-grained Uinta Formation shales in the eastern part of the Uinta basin, contrasting with the red sandstone, shale, siltstone, conglomerate, and limestone beds in the west (Stokes, 1986). Today, fluvial rocks of the middle Eocene Uinta Formation (Uintan NALMA) are preserved above Green River Formation strata only in the northern part of the Uinta basin. The late-middle Eocene (Duchesnean NALMA) Duchesne River Formation overlies rocks of the Uinta Formation in the north-central and western parts of the Uinta basin, and the Cretaceous-age Mesaverde Group on Asphalt Ridge. Both the Uinta and Duchesne River formations were deposited in river channels and on adjacent floodplains (Stokes, 1986) and in river deltas (Townsend, 2004). Riparian environments were forested, and the rivers and streams flowed through a savannatype environment with some swampy and marshy areas and forested highlands (Hamblin, 1987). Uinta Formation The Uinta Formation is exposed at the southern base of the Uinta Mountains in an ~80 mi (~130 km) long, and rather narrow band which can be traced as far westward as Duchesne near the Wasatch Mountain Range and as far eastward as the UtahColorado state line (Bryant et al., 1989; Hamblin et al., 1999; Peterson and Kay, 1931). It overlies the lacustrine Green River Formation, with an interfingering contact between the two formations that reflects their closely related depositional history (Hintze, 2005; Ryder et al., 1976). The Uinta Formation is ~1298 m (4257 ft) thick (including subsurface and exposed rock) and contains bitumen in the form of gilsonite veins (Cashion, 1986; Hintze, 2005; Sprinkel, 2007). The Uinta Formation was deposited

Middle Eocene rock units in the Bridger and Uinta Basins under fluvial conditions as the lake depositing the Green River Formation was receding (Ryder et al., 1976). History of Paleontological Investigations O.C. Marsh led the first paleontological expedition to the Uinta basin in 1870 as part of a Yale College Scientific Expedition (Betts, 1871; Marsh, 1875a, 1875b). In an unexplored area between the Green and White rivers, he and his party collected numerous fossil mammals, the beginnings of an assemblage that would later define the Uintan North American Land Mammal Age (NALMA), a time period that spans ~40–46.5 Ma (Murphey and Evanoff, 2007; Prothero, 1996; Wood et al., 1941). In the 1880s, Francis Speir of Princeton University collected in this same area and the mammalian fossils he collected were studied by W.B. Scott and H.F. Osborn, whose efforts produced the first comprehensive publications on mammals from the Uinta Formation (Rasmussen et al., 1999a; Scott and Osborn, 1887, 1890). At the turn of the century, the Carnegie Museum sent expeditions to the Uinta basin at different times led by O.A. Peterson and then Earl Douglass (Douglass, 1914; Peterson, 1919). Peterson (1919) initially described many of the medium and smaller-sized mammals from the Uinta Formation, many now recognized as index taxa for the Uintan NALMA and therefore crucial for understanding Uintan biochronology on both the regional and continental scales (Walsh, 1996). After these initial expeditions, museums across North America sent parties to collect small samples of Uintan mammals, the most notable of these being the Smithsonian Institution. In 1993, an expedition from Washington University in St. Louis began what would be a 15+ year project (still ongoing) in the Uinta and Duchesne River Formations, with the goal of collecting small mammals (Rasmussen et al., 1999a; Townsend et al., 2006). These latest expeditions have amassed one of the largest assemblages of Uintan mammals from the Rocky Mountain region (Townsend, 2004). Stratigraphy While the Uinta Formation has been formally divided into a lower Wagonhound Member and an upper Myton Member, many geologists and paleontologists continue to utilize the informal tripartite scheme in which the formation was initially subdivided: Uinta A, B and C, from lowest to highest (Prothero, 1996; Sprinkel, 2007; Wood et al., 1941). The tripartite division of the formation is a modification of Osborn’s (1929) stratigraphic nomenclature: Uinta A, Uinta B1, Uinta B2, and Uinta C (Peterson and Kay, 1931; Prothero, 1996; Sprinkel, 2007; Townsend et al., 2006; Walsh, 1996). The lowermost unit, Uinta A, intertongues with the underlying Parachute Creek Member of the Green River Formation and is comprised of resistant fine-grained sandstones that are yellowishbrown and yellowish-gray (Sprinkel, 2007). These beds are medium to massive sandstones that weather to cliffs, ledges, and softer gray slopes and have been measured to be ~226 m

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(740 ft) in thickness (Cashion, 1974). Within in Uinta A, there is a 1.8 m (6 ft) thick tuffaceous bed (“a” tuff), and the base of this bed occurs 56 m (185 ft) below the top of the Uinta A (Cashion, 1974). Uinta B is composed of light-gray, light greenish-gray, light-brown, light-purple mudstones and claystones forming non-resistant slopes (Sprinkel, 2007). These are interbedded with green-gray, yellow and brown fine-grained sandstones that form thin resistant ledges (Sprinkel, 2007; Townsend et al., 2006). There are prominent gilsonite veins in this unit (Cashion, 1986). In the eastern part of the basin, the Uinta B is capped by a massive fine- to coarse-grained gray and brown sandstone bed that weathers to dark brown, and was named the “Amynodon” sandstone (Riggs, 1912). This massive sandstone unit is over one mile in length, and was originally defined as the boundary between the Uinta B and Uinta C horizons (Riggs, 1912), but is now known to be local in extent. Uinta C is distinctive in that the lower intervals and upper intervals are strikingly different in terms of rock type and coloration (Townsend et al., 2006). The lower part of Uinta C is comprised of light-gray, light greenish gray, and light-brown mudstone and claystone that interbed with fine-grain sandstone beds, which are also greenish-gray or commonly yellow or brown (Sprinkel, 2007, Townsend et al., 2006). The upper part is composed of mainly deep red orange, red, dark-brown, grayish-purple and yellow mudstone and claystone with thin gray and green fine grain sandstone beds that form thin ledges (Sprinkel, 2007; Townsend et al., 2006). The stratigraphy of the Uinta Formation has been less well studied than that of the Green River Formation, and most of the stratigraphic work has been undertaken by paleontologists. The early work that produced the initial description of the formation has been confused by changes in terminology describing the lithologic horizons and by shifting boundaries of the units that divide these horizons (Cashion, 1986; Osborn, 1929; Prothero, 1996; Riggs, 1912; Stagner, 1941; Townsend, et al., 2006). Osborn (1895) cited Peterson as first dividing the formation into informal units Uinta A, B, and C, and all units were reported to be fossiliferous. All early fossil collections have these horizon designations assigned as stratigraphic data (Prothero, 1996). Osborn (1929) removed the top 400 ft of Peterson’s Uinta A, and called it Uinta B1, leaving the remaining sandstones of Uinta A without any known record of fossil vertebrates. Additionally, he designated the “Amynodon” sandstone of Riggs (1912) as the boundary unit between Uinta B (now B2) and Uinta C, further altering Peterson’s initial scheme. Osborn’s (1929) Uinta B1 had an upper limit bound by another massive unit, the “Metarhinus” sandstone. The massive sandstone units of Uinta B1 are not laterally extensive, and a red shale and siltstone unit described by Cashion (1974) might be a more appropriate boundary unit. The various changes in terminology and marker units have caused significant confusion in museum locality data, and have adversely affected the accuracy of the stratigraphic ranges of numerous Uintan taxa. Prothero (1996) constructed a magnetostratigraphic section correlating the various historical Uinta Formation localities. With

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the local ranges of various mammalian taxa in disorder due to earlier stratigraphic nomenclatural changes, Prothero’s (1996) work was an essential first step in clarifying and correlating localities across the exposed formation in the Uinta basin. Townsend et al. (2006) described a series of stratigraphic sections through the upper intervals of Osborn’s (1929) Uinta B2 and Uinta C, up to the contact with the overlying Duchesne River Formation (Fig. 14A). These authors suggested a new boundary between the Uinta B–Uinta C units based on a significant lithologic and color change. The suggested boundary is ~73 m higher than the “Amynodon” sandstone of Osborn (1929). Sprinkel (2007) mapped the boundary between the Uinta C and the Duchesne River Formation higher in the section, and because it is associated with such a gradual lithology change in the eastern part of the basin, he recognized and mapped a transition zone between the formations. He mapped the Uinta B-C boundary using a distinctive and reportedly more laterally persistent greenish mudstone bed that occurs just below the “Amynodon” sandstone. Stratigraphic work continues in conjunction with the current paleontological work. Paleoenvironmental History, Paleontology and Biochronology The Eocene epoch is the time when many modern orders of mammals first appear in the fossil record (e.g., primates, bats, artiodactyls, perissodactyls, rabbits,). Prior to the Eocene, mammalian faunas were composed of archaic forms that had diversified just after the demise of the dinosaurs at the CretaceousTertiary boundary. During the early Eocene, global climates were exceedingly warm and produced an almost pole-to-pole tropical greenhouse (Zachos et al., 2001). The Uintan NALMA and the final intervals of the preceding Bridgerian, mark the beginning of the end of this global greenhouse, and this is reflected in the ecological diversity of the mammals that lived during this time (Townsend et al., 2010; Woodburne et al., 2009a). Tropical arboreal forms, typical of the early Eocene greenhouse, began to drop off, and mammals that were more adapted to subtropical, even temperate conditions and habitats began to appear (Townsend et al., 2010). The Uintan NALMA also marks a key transition in the evolutionary history of North American mammalian faunas as ~31% of modern mammalian taxonomic families (ancestors of canids, camelids, felids, and some modern rodents) appeared in the fossil record (Black and Dawson, 1966). The Uinta Formation fossils that form the basis of the Uintan NALMA are from Uinta B1, B2, and C units. The fauna from Uinta B1 comes mainly from the Wagonhound-Bonanza area, an early locality (Well No. 2), and the Willow Creek-Ouray area. The assemblage from this unit is small and includes Achaenodon uintensis, Protoreodon parvus, Oromeryx plicatus, Harpagolestes, Eomoropus amarorum, Hyrachyus eximius, a species of Triplopus, and a diverse group of brontotheres and uintatheres (see Appendix A) (Gunnell et al., 2009; Prothero, 1996). These taxa are considered an early Uintan assemblage corresponding to biochron Ui2 of Gunnell et al. (2009).

The Uinta B2 fauna is incredibly diverse (see Appendix A) and constitutes the bulk of the assemblage that is considered early Uintan, and corresponds to biochron Ui2 of Gunnell et al. (2009). Sediments that have yielded faunas from the Uinta B2 unit include White River Pocket and areas around Kennedy’s Hole. Recent fossil collecting efforts have made it possible to formally define biochronologic units and stratotype sections for the Uinta Formation (Gunnell et al., 2009). The beginning of Ui2 is defined by 26 index taxa (outlined in Gunnell et al., 2009). Numerous taxa range through Ui2, but only a few are unique to Ui2 and these are from Uinta B2 beds: Amynodon reedi, Eobasileus cornutus; Eomoropus amororum, Hyrachyus eximius, Limnocyon potens, Mesonyx obtusidens, Metarhinus fluviatalis, Metarhinus sp., Ourayia sp. nov., Pantolestes longicaudus, Paramys sp., Sciuravus popi, Uintatherium anceps (Gunnell et al., 2009). Fossil mammals found in Uinta C sediments are typically considered to be late Uintan and correlate to biochron Ui3 of Gunnell et al. (2009). Strata that have yielded late Uintan faunas include Myton Pocket, Kennedy’s Hole, Leota Ranch, Devil’s Playground, Leland Bench Draw, and Antelope Draw. The Uinta C fauna is equally as diverse as the Uinta B2 fauna. Twenty-four taxa define the beginning of Ui3 including Auxontodon, Colodon, Duchesneodus, Eosictis, Janimus, Mytonius, Mytonomeryx, Pentacemylus, Procyonodictis, Protadjiduamo, Protitanotherium, and Metatelmatherium. One of the striking differences between Ui2 and Ui3 faunas is that the younger assemblage is dominated by artiodactyls. Therefore, this biochron was designated the “Pentacemylus progressus Interval Zone” reflecting the abundance of this particular taxon (Gunnell et al., 2009). While the taxonomic composition of Ui3 is quite different than Ui2, only a few genera make their first appearance during this interval (Gunnell et al., 2009). Current fieldwork is focused on collecting the upper intervals of the Uinta C unit via excavation and screenwashing in order to increase the sample of small mammals. Uinta Formation Field Trip Stops Traveling south across the Uinta basin via State Route (SR) 45, we will cross both the Green and White rivers and make our way to the base of the Uinta Formation just south of Wagon Hound Canyon. Note that the driving distance from the Utah Field House Museum in Vernal to the first field trip stop is 42.6 mi. Our examination of the Uinta Formation will commence at the base of the formation where it overlies strata of the Parachute Creek Member of the Green River Formation along the White River. From there we will travel north across Wagon Hound Canyon to examine rocks of the Uinta A and interbedded Green River Formation. Continuing north, we will examine the green siltstones and brown sandstones typical of Uinta B1. From here we will be able to see the typical red and gray beds defining Uinta B2 and the orange-red sediments that typify the Brennan Basin Member of the Duchesne River Formation from a distance. We will spend most of the day exploring typical Uinta B2 localities in the Coyote Wash area and will then

Middle Eocene rock units in the Bridger and Uinta Basins

Figure 14. (A) Stratigraphic section from Townsend et al. (2006) indicating the level of the Amynodon sandstone of Osborn (1929), the base of which forms the traditional Uinta B2–Uinta C boundary, and the Uinta B2-C boundary at Devil’s Playground suggested by Townsend et al. (2006). (B) View of massive cliff-forming Uinta A sandstones overlying and interfingering with Parachute Creek Member of the Green River Formation. (C) Measuring one of the massive sandstone units in Uinta B. (D) Collecting at WU-18, a typical Uinta B2 locality. (E) Sandstone channels associated with the WU110 locality complex.

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drive to the Red Wash area where the transition to Uinta C rocks is well exposed. From Red Wash, we will make our way up section a high Uinta C fossil locality near the contact with the overlying Duchesne River Formation. Our final Uinta Formation stop is located to the southwest of this location at a classic Uinta Formation fossil locality, White River Pocket. Stop 1. Uinta Formation–Green River Formation Contact along the White River at Wagon Hound Canyon (0.0 mi; 0.0 cumulative; SR 45 highway milepost 0.8) Peterson in Osborn (1895) described the contact between the Uinta Formation and the Green River Formation as a series of “hard brown sandstones” that conformably overlie the Green River shales. Looking north across the White River, the Parachute Creek Member of the Green River Formation is exposed as sloping yellow-brown, tan, and green-gray mudstone and claystone directly underlying the massive sandstone cliffs of Uinta A (Fig. 14B). Looking to the east along the White River, the Green River shales are exposed as dark-brown bands within the member. The “hard brown sandstones” of the Uinta Formation are markedly different from the underlying Green River Formation in that they form massive resistant cliffs of yellow brown sandstone. There are few reported identifiable vertebrate fossils from the Uinta A with stratigraphically reliable locality data, and none collected thus far are biostratigraphically diagnostic. The Uinta A may be earliest Uintan (Ui1a) in part (or possibly even Br3 in part), and therefore could be partially equivalent with the Turtle Bluff Member of the Bridger Formation (see Murphey and Daitch, 2007). Head north on SR 45 and take a left at 1.7 mi (milepost 2.4) and turn west onto a graded dirt road. Bear right at the bottom of the hill 0.1 mi ahead (cumulative 1.8 mi from Stop 1), and turn onto a less well maintained road. Pull off after 0.1 mi (cumulative 1.9 mi from Stop 1). Stop 2. Uinta A and Cashion’s “a” Tuff (1.9 mi from Stop 1; 1.9 cumulative mi) Driving north on Highway 45 through Wagon Hound Canyon we turn west onto a graded dirt road. Here the sloping yellow-gray siltstones and cliff forming sandstones of Uinta A are well exposed. Cashion’s (1974) “a” tuff is exposed on the slope and is most visible just below the cliffs to the north. This tuff has not yet been dated because it is so crystal poor. The boundary between the Uinta A and Uinta B is mapped as occurring half way up the cliff above the tuff. Beds of lacustrine shale are interbedded with the lower part of the Uinta A sequence in this general area, reflecting the complex transition from lacustrine to fluvially dominated depositional environments. Fragments of fossilized bone and wood have been found in some of the side canyons in this area. Optional Stop Peterson’s 1924 Dolichorhinus Quarry. Continue southwest from Stop 2. After passing a small gilsonite mine the road will

bend to the northwest. After another 0.15 mi, there is a fork in the road. Take the right fork to the top of the hill to your north. At ~1.8 mi from Stop 2, the spoils pile from Peterson’s 1924 Dolichorhinus Quarry will be visible 0.2 mi to the east of the road. This historic quarry site was re-located by Vernal paleontologists Evan Hall and Sue Ann Bilbey. Return to Stop 2, then get back on SR 45 and head north and northeast past the Bonanza gilsonite mining operation. Park on the east shoulder of the highway (3.4 mi from Stop 2). Stop 3. Uinta B1 and Metarhinus Sandstone (3.4 mi from Stop 2; 5.3 cumulative mi; SR 45 highway milepost 5.6) This unit is composed of a massive yellow-gray sandstone bed underlain by greenish-gray siltstone (Fig. 14C). The Metarhinus sandstone is not a widespread unit. It is the boundary between Osborn’s (1929) Uinta B1 and B2. Off to the west both the Bonanza and Independent gilsonite veins are visible. The most common vertebrate remains found here are turtle and crocodile. Mammals that have been recovered include mostly brontotheres and other large Uintan mammals. Continue northwest on SR 45. At 3.6 mi from Stop 3 (SR 45 highway milepost 9.3), turn northeast at the intersection with Uinta County Road 3150. The dark-brown, bench supporting, sandstone bed to your east and west is the “Amynodon sandstone,” which has traditionally formed the boundary between the Uinta B and Uinta C. Continue down Uinta County Road 3150 for 4 mi. After crossing the railroad tracks, there is a dirt road turning south that crosses over the tracks. Head south down this dirt road. At ~5.25 mi from Stop 3, there will be a two-track road on the right, heading west. This two-track goes across Quaternary sediments with Uinta Formation outcrops to the north. The two-track road will veer to the right, toward the outcrops. This is the region of the WU-18 locality. Stop 4 is located 6.4 mi from the junction of SR 45 and Uinta County Road 3150, and 10.49 mi from Stop 3. Stop 4. Uinta B2 and WU-18 (10.49 mi; 15.79 cumulative mi) Driving into Coyote Wash, sediments characteristic of Uinta B2 lithologies can be observed: greenish-gray, yellow, and light-pink to red mudstone and claystone. The WU-18 locality or “Gnat-Out-Of-Hell,” is a typical Uinta B2 locality in terms of the lithology (Fig. 14D). WU-18 is the single most productive Uinta B2 locality currently known. The main productive level is a dark-brown sandy mudstone that has produced numerous mammals, including the early primate Chipetaia and large bodied amynodont rhinocerotoids and brontotheres (Rasmussen, 1996; Rasmussen et al., 1999a). Coprolites with small mammal bones from WU-18 indicate that this locality may have been partially accumulated by a diurnal raptor (Thornton and Rasmussen, 2001). Stratigraphically above and to the northeast of WU-18 are the “Amynodon sandstones,” the classic boundary between Uinta B2 and Uinta C (Osborn, 1929). This massive sandstone unit sometimes yields large-bodied ungulates, such as its rhinocerotoid namesake, Amynodon, and brontotheres. Screen washing

Middle Eocene rock units in the Bridger and Uinta Basins at WU-18 has resulted in the recovery of multiple bird bones, increased samples of primates and small rodents. Return to SR 45 and head north. Pull over carefully as close to the guardrails as possible at SR 45 highway milepost 11.6, which is 1.8 mi from Stop 4. Stop 5. Devil’s Playground Red Beds (1.8 mi from Stop 4; 17.59 cumulative mi; SR 45 highway milepost 11.6) To the east there is a prominent red bed that Townsend et al. (2006) consider to be the boundary between the Uinta C and the Brennan Basin Member of the Duchesne River Formation. Heading south on SR 45, about ¾ of a mile, make a sharp right and travel a mile and veer to the right at ~1 mi from the turn off from SR 45. Travel for another 1¼ mi to a well pad and park. Stop 6. Uinta C WU-110 Locality Complex (2.25 mi; 19.84 cumulative mi) The WU-110 locality complex spans the Uinta B2–Uinta C boundary as defined by Townsend et al. (2006). The primary lithology in the area consists of gray-green, tan, and yellow-brown mudstones. A distinctive feature of this complex includes large dark-brown fine- to coarse-grained channel sandstone beds (Fig. 14E). Numerous fossils have been discovered in a dark-brown siltstone layer that weathers pink at WU-110. These fossils include the primate Ourayia, the artiodactyls Protoreodon and Leptotragulus, as well as numerous rodents. The remaining localities in the area WU-115, 116, 117, have also been quite productive. In addition to more small mammal dentaries, large-bodied perissodactyls, carnivore postcrania, and a specimen of Simidectes have been recovered. Return to SR 45. Turn right and travel ¾ of a mile to the first dirt road exit on the left, this is the Sand Ridge Road. Take a right. Follow Sand Ridge Road west, and at mile 7 from Stop 6, you will come to a road that veers left. Do not turn here; continue west. After mile 9, the road will turn south, and at mile 10 you will encounter a fork in the road. Make a right here, and in another ½ mi there will be another fork where you will take a right again. Continue on this road for 3.5 mi. Close to mile 14, you will see a road veering left, do not turn here, continue straight ahead. At approximately mile 15, there is a road that veers right (north). Turn here. At the 16.5 mi mark, pull off the road and park. Looking due west, you will see a large gray and orange outcrop in the distance, this is the WU-22 locality complex. Stop 7. Uinta C and WU-22 Locality Complex (16.69 mi; 36.53 cumulative mi) Note the Deseret Power Plant to the northeast and a large bench directly north of the locality complex with strata consisting of highly variegated red, purple, gray, and orange sediments. These claystone and mudstone beds surround the main drainage of this region of the basin, Red Wash. Fossils at WU-22 are generally found in gray, tan, yellow, and green-gray mudstone beds. Typically, as with other localities, fossils here are found directly on the surface, where they have eroded out of the surrounding

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sediments and are highly fragmented. Small-scale excavations in this locality complex have recovered a number of articulated skeletons, including the skull and postcrania of the tiny deerlike Protoreodon and a skull, forelimb, and vertebral column of medium sized rodent. From Stop 7, turn left, heading northeast for ~2 mi until reaching an intersecting road, and turn left again. After ¼ mi, you will see an intersection of four roads. Take the road to your immediate left, which will lead west along a pipeline road. Stay on this road for the next 5–6 mi. This is a main transport road in the oil fields. After mile 7, this road will become Glen Bench Road. At about mile 7¾, the road will jog to the right and you will see an intersecting road. Turn left and begin the steep climb up to the top of Glen Bench. There is a sharp turn at about mile 9¼. After this turn, continue along the road until you come across an intersecting road just past mile 10. Cross this intersecting road and follow it down the northwest face of Glen Bench into Antelope Draw. There will be a three-way intersection at about the 11¾ mi mark. Veer left, heading northwest. At mile 13, take a sharp left, driving for more than a ½ mi. Make a hard right heading due north, this is another major transport road. At ~14½ mi, you will come to two roads that veer off to the right, one heading northeast, the other southeast. Take the northeast road. After mile 15, the road will loop around a large canyon. Continue along a series of tall escarpments, and the road will make another loop after mile 17. At this point, find a place off of the main road to park. Looking due east, you will see the WU-210 locality complex. Stop 8. Uinta C, WU-210, Contact with Duchesne River Formation (18.71 mi; 55.24 cumulative mi) Nearing the top of the Uinta Formation, the sediments are brightly pigmented as they weather deep brown, bright orange, and various shades of red. In the WU-210 area, sloping red, gray, and tan mudstone beds are interbedded with large, ledgeforming blocky yellow-brown sandstone beds. At this level in the formation, the Duchesne River and Uinta formations are locally interbedded. Conglomeratic resistant sandstone channels typical of the Brennan Basin member of the Duchesene River Formation can be viewed at the top of WU-210. This locality and others surrounding it are highly productive and have yielded numerous perissodactyls including early tapiroids, such as Isectolophus, and almost complete skeletons of large rodents such as Pseudotomus. From WU-210, head north toward the loop in the road and make your way back to the main road from which we approached this complex of localities. Take a right heading north going toward the top of Deadman Bench. This road will intersect with the main Deadman Bench road that trends east-west. Take a right on this main road and head back to SR 45 (6.6 mi). Head southeast on SR 45 for 5.9 mi and turn northeast onto a short graded road at SR 45 highway milepost 14.0. There is an old information kiosk and small wind turbine located ~300 ft east of the highway at this location.

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Stop 9. Uinta Formation—Duchesne River Formation lower Transition Zone Contact (16.2 mi from Stop 8; 71.44 cumulative mi; SR 45 highway milepost 14.0) This is the approximate stratigraphic location of the base of the transition zone between the Uinta C (Myton Member) and the Brennan Basin Member of the Duchesne River Formation as mapped by Sprinkel (2007). Drive northwest along SR 45 for 0.7 mi until you come to an intersection with the paved road which leads to the Deseret Power Plant heading southwest and stop. Stop 10. Uinta Formation–Duchesne River Formation Upper Transition Zone Contact (0.7 mi from Stop 9; 72.14 cumulative mi; SR 45 highway milepost 14.7) Looking east you can view the approximate stratigraphic location of the top of the transition zone between the Uinta C (Myton Member) and the Brennan Basin Member of the Duchesne River Formation as mapped by Sprinkel (2007). Continue northwest on SR 45 and pull over at SR 45 highway milepost 31.5, 16.7 mi from Stop 10. Stop 11. View of Asphalt Ridge (16.7 mi from Stop 10; 88.84 cumulative mi; SR 45 highway milepost 31.5) For the last 16.7 mi, you’ve been driving through rocks of the Brennan Basin Member of the Duchesne River Formation. This stop affords an excellent view of Asphalt Ridge, a geologically interesting structure composed of uplifted and tilted strata of the Brennan Basin Member of the Duchesne River Formation unconformably overlying uplifted and tilted strata of the late Cretaceous Mesa Verde Group. The structural geology of Asphalt Ridge documents different episodes of uplift and erosion related to Uinta Mountain tectonics. The final Uinta Formation field trip stop is a visit to the classic fossil locality White River Pocket. Detailed directions to this field trip stop are not included in this field trip guide because there are numerous ways to get there. As a result, the mileage to this field trip stop is not connected to the preceding field trip stops. To get to White River Pocket, travel to the point at which SR 88 crosses the White River Bridge south of the hamlet of Ouray. Stop 12. White River Pocket Head South on SR 88 for 0.2 mi. Pull over onto the shoulder. The badland hill and surrounding area to your east is a classic fossil locality called White River Pocket. The fossil assemblage from White River Pocket provided the basis for early Uintan faunas and are significant because a large number of small mammal taxa were found there. Drive north on SR 88 for 11.7 mi. Just past the Pelican Lake Café, turn left onto Uinta County Route 2762. Drive west for 4.7 mi and park on the north shoulder. To your right are spectacular cliffs that are historically known to paleontologists as “Randlett Point,” with excellent exposures of the sequence that contains the conformable boundary between the uppermost Uinta

C (Myton Member) and the Brennan Basin Member of the Duchesne River Formation, type sequence of the “Randlett Horizon.” For further explanation, see Stop 1 of the Duchesne River Formation field trip below. START OF DUCHESNE RIVER FORMATION TRIP Duchesne River Formation The Duchesne River Formation occurs only in the northern and western parts of the Uinta basin in Utah, where it is widespread. Its definition and stratigraphic history are related to the underlying and apparently conformable Uinta Formation. In the northern part of its distribution along the Uinta Mountain front, it unconformably overlies rocks of Triassic and Jurassic age. Interestingly, the Duchesne River Formation contains tar sand in the Asphalt Ridge area (U.S. Geological Survey, 1980). The Duchesne River Formation was deposited under mostly fluvial conditions but includes some minor lacustrine deposits. Andersen and Picard (1972) proposed the presently accepted stratigraphy and described the following members in ascending stratigraphic sequence: Brennan Basin, Dry Gulch Creek, Lapoint, and Starr Flat. The Brennan Basin Member is composed of soft to moderately resistant, light-to-medium red, light-gray, light-brown, yellow, and tan ledgy sandstone, mudstone, conglomerate, shale, and siltstone, with a maximum thickness of ~600 m (1970 ft) south of Vernal. This member thins significantly to the east and west. The Dry Gulch Creek Member consists of soft to moderately resistant, light-to-medium gray, medium red, purplish gray, and yellow sandstone, mudstone, shale, and conglomerate. It is ~149 m (490 ft) thick. The Lapoint Member consists of mostly soft, light-red, tan, and yellow sandstone, siltstone, and mudstone with minor amounts of conglomerate. It contains diagnostic beds of light-gray to medium-gray or bluish-gray bentonite and ranges in thickness from ~119–299 m (390–980 ft). The Lapoint ash, which forms the base of the Lapoint Member (where most of the vertebrate fossils have been discovered), has been dated at 39.74 Ma ± 0.07 Ma (Prothero and Swisher, 1992). The Starr Flat Member consists of moderately resistant, light-to-medium red and tan sandstone, mudstone, with significant red and gray conglomerate, and with a thickness of ~149 m (490 ft) (Rowley et al. 1985). It outcrops spottily along the southern flank of the Uinta Mountains. The history of stratigraphic nomenclature for the Duchesne River Formation is somewhat confusing (Table 5). Clarence King (1878) named the Uinta Group for what we now call the Uinta and Duchesne River formations. Peterson, in Osborn (1895), proposed that the Uinta Group be subdivided into A, B, C horizons, with the C horizon being equivalent in part to what we now know as Duchesne River Formation and equivalent in part to the upper Uinta Group. Douglass (1914) suggested that the name Uinta Group be replaced by the name Uinta Tertiary to avoid confusion with the Precambrian Uinta Mountain Group.

Brennan Basin Member

Dry Gulch Creek Member

Minor bentonite member

Lapoint Member

Halfway horizon Duchesne Formation Uinta Tertiary (in part) Uinta Group (in part)

C horizon (in part) = upper Uinta Group

Upper Uinta redbeds separated from underlying Uinta Tertiary

Randlett horizon

Halfway horizon

Lapoint horizon Lapoint horizon

(not studied)

Major bentonite member

(not studied) (not studied)

Gazin (1955) Kay (1934) Scott in Peterson (1931) Douglass (1914) King (1878)

Peterson in Osborn (1895)

Peterson and Kay (1931)

Duchesne River Formation

TABLE 5. HISTORY OF DUCHESNE RIVER FORMATION STRATIGRAPHIC NOMENCLATURE

Warner (1963)

Andersen and Picard (1972) Starr Flat Member

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Scott in Peterson (1931c) named the Duchesne Formation, and it was renamed Duchesne River Formation by Kay (1934) because the name Duchesne Formation was preoccupied. Kay (1934) defined the Duchesne River Formation by removing the red beds of Osborn’s (1895) upper Uinta Group, recognizing they have a younger mammalian fossil fauna. He also recognized three biostratigraphic “horizons:” Randlett, Halfway, and Lapoint. Gazin (1955) removed the Randlett horizon from the Duchesne River Formation, recognizing that it contains a fauna that is almost identical to that of the Uinta C. Warner (1963) recognized two members, the Minor Bentonite member being equivalent to the Brennan Basin and Dry Gulch Creek members, and the Major Bentonite Member being equivalent with the Lapoint Member. Andersen and Picard (1972) named the currently recognized members: Brennan Basin, Dry Gulch Creek, Lapoint, and Starr Flat members. With regard to Andersen and Picard’s (1972) stratigraphic terminology, the Brennan Basin Member is equivalent to the Randlett Horizon of Kay (1934) and the lower part of Halfway Horizon, the Dry Gulch Creek Member is equivalent with the upper part of the Halfway horizon, and the Lapoint Member is equivalent to the Lapoint horizon. Fossils and Biochronology of the Duchesne River Formation In comparison with the underlying Uinta Formation and correlative strata of Duchesnean age in coastal southern California, the vertebrate fossil fauna of the type Duchesne River Formation is sparse. However, it is a critically important period in mammalian evolution (Lucas 1992; Rasmussen et al., 1999b; Robinson et al., 2004). Although it is the nominal stratotype for the Duchesnean NALMA (Wood et al.,1941), its validity has been the subject of some controversy among paleontologists (see Lucas, 1992; Wilson, 1978). Nevertheless, the Duchesnean NALMA has now been widely accepted by paleontologists (Rasmussen et al., 1999b; Robinson et al., 2004). Lucas (1992. p. 88) made the observation that the Duchesne River Formation “has either been questioned, abandoned, subdivided, or defended.” Scott (1945) regarded the Duchesne River Formation as Oligocene in age. Simpson (1946) and Gazin (1955) considered it Eocene. Faunally, the importance of the Duchesne River Formation and the Duchesnean NALMA is based on the fact that it records a major faunal replacement in North America, as demonstrated by its large number of first and last occurrences, as well as the small number of genera that are restricted to it (Appendix A; Black and Dawson, 1966; Robinson et al., 2004). Robinson et al. (2004) tentatively assigned Duchesnean first appearances to include Hyaenodon, Duchesneodus, Duchesnehippus intermedius, Amynodontopsis, and Eotylopus. The fauna of the Brennan Basin Member, which was previously assigned to the Randlett horizon and lower part of the halfway horizon of Kay (1934), includes a faunal assemblage that is generally considered to be intermediate between the Uinta C (biochron Ui3, Gunnell et al., 2009) and the Lapoint Member

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of the Duchesne River Formation, although most workers regard it as late Uintan. The Brennan Basin Member outcrops over a fairly large area, and several somewhat productive localities are known. It contains a reasonably diverse assemblage of Uintan artiodactyls including Protoreodon, Pentacemylus, Diplobunops, and Leptotragulus. Perissodactyls include a mix of Uintan and Duchesnean brontotheres, as well as tapirs and rhinos. Specimens of the Uintan lagomorph Mytonolagus and the rodents Pareumys and Mytonomys have also been collected (Rasmussen et al., 1999b). Until recently, only two published specimens had ever been reported from the Dry Gulch Creek Member, and these represented the “Halfway Fauna.” The specimens include the holotype of Duchesnehippus intermedius, a genus of horse that is purportedly intermediate between the Uintan horse Epihippus and the Chadronian horse Mesohippus, and a femur and astragalus tentatively identified as (?) hyracodontid rhinoceros. Walsh and Murphey (2007) discovered a locality that produced a diversity of small mammals in the Dry Gulch Creek Member (67 identifiable specimens). SDSNH Loc. 5939 (“Halfway Hollow One”) is situated roughly in the stratigraphic middle of the Tdr-dgc, at about the same level as the type locality of D. intermedius. Taxa from this locality include Nanodelphys sp., Lipotyphlan indet., Pareumys cf. P. guensburgi, cf. Janimus sp., Metanoiamys sp., Protadjidaumo typus, Protadjidaumo sp. (large), Eomyidae, unidentified gen. and sp., Eomyidae? new gen. and sp., Geomorpha, new gen. and sp., cf. Griphomys sp., cf. Passaliscomys sp., Mytonolagus sp., and Artiodactyla indet. Thus, the number of identifiable mammal specimens from the Dry Gulch Creek Member has been increased dramatically. Fourteen additional mammalian taxa are now known from this member, nine of which are new records for the Duchesne River Formation as a whole. Importantly, Pareumys cf. P. guensbergi, Protadjidaumo typus, and cf. Passaliscomys sp. suggest a Duchesnean age for most or all of the Dry Gulch Creek Member, and that the Uintan-Duchesnean boundary occurs somewhere within the upper part of the Brennan Basin Member or the lower part of the Dry Gulch Creek Member. Additional stratigraphically controlled screen washing efforts are needed to determine the precise stratigraphic location of the Uintan-Duchesnean boundary in the Uinta basin. The mammalian fauna of the Lapoint Member is considered the type Duchesnean fauna, although misidentifications, data loss, and mistakes made by earlier workers have resulted in considerable confusion regarding what was actually collected at this horizon. Most of these problems seem to have been resolved by the efforts of Rasmussen et al. (1999b). The Lapoint fauna contains diverse assemblages of artiodactyls such as Protereodon, Agriochoerus, Simimeryx, and Brachyops, and perissodactyls such as Colodon, Hyracodon, Duchesnehippus, and Duchesneodus. Also present are carnivores, the mesonychid Hessolestes, the hyaenodontan Hyaenodon, the lipotyphlan Centetodon, and the rodents Pareumys and Protadjidaumo. No published accounts of fossils have been reported from the stratigraphically highest Starr Flat Member. However, Mark

Roeder and Paul Murphey of the San Diego Natural History Museum discovered a highly fragmented mammal tooth that was tentatively identified as a partial upper molar of a large camelid from near the base of the member (Paul Murphey, 2008, unpublished field notes). Although the Lapoint and Brennan Basin members of the Duchesne River Formation are the most fossiliferous, fossils are generally scarce throughout the formation. Any new discoveries would be highly significant, especially from the Dry Gulch Creek and Starr Flat members. It is also noteworthy that reptiles in general are poorly represented from the Duchesne River Formation. Much work remains to be done in order to document the faunal changes within the Duchesne River and ascertain the position of the Uintan-Duchesnean boundary. Duchesne River Formation Field Trip Stops Our exploration of the Duchesne River Formation begins at Randlett Point where the contact between the Myton Member of the Uinta Formation (informal member C of Uinta Formation) and the overlying Brennan Basin Member of the Duchesne River Formation is well exposed and relatively easy to discern. From there we travel north through strata of the Brennan Basin Member in Halfway Hollow, and then examine exposures of the Dry Gulch Creek Member and discuss the paleontology of this sparsely fossiliferous unit that appears to contain the boundary between the Uintan and Duchesnean NALMA’s although the precise stratigraphic position of the boundary has not yet been determined. From a distance, we then view the Lapoint Member and overlying Starr Flat Member in Halfway Hollow. Heading east along the Lapoint highway to Twelvemile Wash, our last stop is at the site of the famous Carnegie Museum Titanothere Quarry just above the base of the Lapoint Member, as well as the Lapoint ash at quarry level. Stop 1. Randlett Point (0.0 mi; cumulative 0.0 mi) Peterson and Kay (1931) defined the base of the Duchesne River Formation at this location. The boundary between the Uinta C (Myton Member) and overlying Brennan Basin Member in this part of the Uinta basin is easier to discern than further to the east in the type area of the Uinta Formation in the Coyote basin. Here the contact is marked by the change from slope-forming pink, purple, light-gray, and red variegated beds of mudstone and siltstone to orange, reddishbrown and gray ledge forming siltstone and sandstone beds of the overlying Brennan Basin Member (Fig. 15A). The Carnegie Museum’s Randlett Quarry is ~1 mile to the north of this location. With the exception of the large Duchesnean brontothere Duchesneodus, the fauna of the Brennan Basin Member resembles that of the Uinta C, and the fauna of the lower part of the Brennan Basin Member is regarded by most workers to be Uintan in age. Drive east on the Ouray-Randlett Road (Uinta County Road 2762) for 4.7 mi to the junction of State Route (SR).

Middle Eocene rock units in the Bridger and Uinta Basins Stop 2. Brennan Basin Member–Dry Gulch Creek Member Boundary (19.0 mi from Stop 1; cumulative 35.6 mi) Andersen and Picard (1972) defined this boundary as the base of the lowest bed of fine-grained rock overlying the highest resistant sandstone of the Brennan Basin Member (Fig. 15B). This contact was mapped by Rowley (unpublished field maps), Rowley et al. (1985), and Sprinkel (2007) as the sandstone bed visible at this stop. This tan sandstone and conglomerate is highly variable in thickness, and crosses Halfway Hollow Road at UTM Zone 12, 609901 mE, 4470184 mN (NAD 27). A widespread but thin 0.5 m (1.6 ft)

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thick dark-red indurated fine-grained sandstone, interpreted by Andersen and Picard (1972) as a paleosol, can be seen ~18 m (60 ft) above the top of the sandstone. Fossil snails and bone fragments have been obtained from screenwashing sediments collected from between the boundary sandstone bed and red paleosol ~0.25 mi to the southwest of this location. However, no identifiable mammal fossils have been found to date in these samples. Continue driving north on Halfway Hollow Road for 2.4 mi. You will stop to the east of a large evaporation pond with 5 tank batteries located along its southern edge.

Figure 15. (A) The contact between the Uinta C (Tu-c) and Brennan Basin Member of the Duchesne River Formation (Tdr-bb) at Randlett Point, Uintah County, Utah. (B) View looking northeast at the sandstone bed (Ss) that divides the Brennan Basin Member of the Duchesne River Formation (Tdr-bb) from the overlying Dry Gulch Creek Member (Tdr-dgc) in Halfway Hollow, Uintah County, Utah. (C) View looking west at the Carnegie Museum Teleodus Quarry (TQ), the lower bentonite bed (LB) that marks the boundary between the Dry Gulch Creek Member (Tdr-dgc) of the Duchesne River Formation and the overlying Lapoint Member (Tdrlp), and the Lapoint tuff (LpT), Uintah County, Utah.

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Stop 3. Dry Gulch Creek Member Fauna (2.4 mi from Stop 2; cumulative 38.0 mi) This part of Halfway Hollow has produced the only known mammal fossils from the Dry Gulch Creek Member. A relatively fossil-rich micromammal quarry was discovered just to the east of the road (Halfway Hollow One, SDSNH Loc. 5939) in 2006, and is discussed above. These fossils are currently being formally described, and it is likely that additional screen washing of sediments from this locality will yield additional mammal fossils. Peterson (1931c, p. 70–71) described the type specimen of Epihippus (Duchesnehippus) intermedius (along with specimens he tentatively identified as a partial femur and astragalus of a hyracodontid rhinoceros) as coming from “the west side of Halfway Hollow, from a red clay about 8 feet thick, overlain by a massive brown sand and underlain by a reddish nodular clay associated with astragalus and horse jaw.” The specimens were collected by John Clark on 21 July 1931 (Alan Tabrum, Carnegie Museum of Natural History, 2006, written commun.). Continue driving north/northeast on Halfway Hollow Road for 0.9 mi until you reach SR 121 (Lapoint Highway). Stop just beyond the cattle guard. Stop 4. View of Lapoint and Starr Flat Members (0.9 mi from Stop 3; cumulative 38.9 mi) Looking north from this spot affords an excellent view of the Lapoint Member, the Starr Flatt Member, and the distant Uinta Mountains with 3668 m (12,031 ft) high Leidy Peak on the skyline. The base of the Lapoint Member is defined as the lowest widespread bentonite bed. Thick, laterally persistent bentonite beds are characteristic of this member. Several additional American Museum of Natural History and Carnegie Museum fossil localities are located in this part of Halfway Hollow to the north of the Lapoint highway, although recent efforts to relocate them have proven unsuccessful. The contact between the Lapoint and overlying Starr Flat members is located at the approximate level of the tops of the red cliffs you see in the distance at the north end of Halfway Hollow. However, the top of the Duchesne River Formation is not visible from this location. As discussed above, the Starr Flatt member has yielded tooth fragments from one locality, but no other fos-

sils have been reported. That said, it has yet to be thoroughly prospected. It is possible that the fauna of the Starr Flatt Member is younger than the Duchesnean NALMA. Turn east onto the Lapoint highway and drive for 2.5 mi. Just past SR 121 highway milepost 31, turn north onto a two-track road and drive for 0.7 mi for a total of 3.2 mi from Stop 4. The Carnegie Teleodus Quarry pit is located near the summit of the large hill to the northwest of your location. Stop 5. Carnegie Titanothere Quarry and Lapoint Ash (3.2 mi from Stop 4; cumulative 42.1) This is the site of the famous Carnegie Titanothere Quarry (also known as the Teleodus Quarry and the Duchesneodus Quarry). For many years prior to the discovery of this quarry, there remained a recognized hiatus between the earlier mammalian faunas of the Rocky Mountain intermontane basins and the younger mammalian faunas of the White River badlands of South Dakota and Nebraska. In discovering and working this quarry, J.L. Kay of the Carnegie Museum made the first important paleontological discovery to begin to fill this knowledge gap. Peterson (1931a, 1931b, 1931c) and Peterson and Kay (1931) described the fossils from this locality and from other localities in the then Lapoint horizon. These workers regarded the fauna as “perfectly transitional” (Peterson, 1931c, p. 62). In addition to the brontothere fossils collected from the quarry, bunodont and selenodont artiodactyls, a rodent, an insectivore, hyracodontid rhinos, and a creodont and mesonychid were discovered and described. Note that a recent thorough reinspection of the quarry and surrounding area has failed to yield any additional fossils. The thin bentonite bed visible on the slope below the quarry is considered to be the base of the Lapoint Member by Andersen and Picard (1972). However, the Lapoint ash (39.74 Ma ± 0.07 Ma, Prothero and Swisher, 1992) is actually the thicker, higher bentonite bed that is cut by the red sandstone channel in which the Carnegie Quarry fossils were preserved (Fig. 15C). Therefore, the Carnegie Titanothere Quarry is located stratigraphically just above the base of the Lapoint Member and is slightly younger than the Lapoint tuff that was cut by the red sandstone channel. End of Uinta basin field trip.

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APPENDIX A. BIOCHRONOLOGIC RANGES OF BRIDGERIAN, UINTAN, AND DUCHESNEAN MAMMALIAN GENERA Orde r Family Genu s Br1a Br1b Br2 Br3 Ui1a Ui1b Ui2 Ui3 Cetartiodactyla Agriochoeridae Agriochoerus Diplobunops X X Protoreod on X X Camelidae Poebrodon X X Dichobunidae Auxontodon X Diacode xis X B u no m er y x X X Hylomeryx X X Mytonomeryx X Neodiacode xis X Pentacem ylus X Entelodontidae Brachyhyops Dyscritochoerus X X Helohyidae Achaenodon Helohyus X X X X Homacodontidae Antiacodon X X X X X Homacodon X Mesomeryx X Microsus X X X Hypertragulidae Simimeryx Oromeryx X X X Oromerycidae Protylopus X X X Protoceratidae Leptoreodon X X X Leptotragulu s X X X Poabrom ylus Carnivora

Miacidae

Viverravidae

Miacis Miocyon Oodectes Procyonodi ctis Tapocyon Uintacyon Vulpavus Didymictis Viverravus

X

X

X

X

X

X

X X X X

X X

X X

X

X

X

X X X

Cimolesta

Apatamyidae Pantolestidae

Apatemys Pantolestes Simidectes

X X

X X

X X

Leptictida

Leptictidae

Palaeictops

X

X

X

Lipotyphla

Sespectidae

Crypholestes Sceno pa g us Centetodon Marsholestes Entomolestes Talpavus Micropternodus Nyctitherium Pontifactor

X X X

X X X

X

X

X X X X X

X

X

Geolabididae Erinaceidae Micropternodontidae Nyctitheriidae

X

X

X

X X

X X

X

X X

X X X

X

X X

X X

X X X

X X

X

X

X X

X X

X

X

X X X

X

X

X

X

Icaronycteridae

Icaronycteris

Hyopsodontidae

Hyopsodus

X

X

X

X

Creodonta

Oxyaenidae Hyaenodontidae

Patriofelis Apataleurus Hyaenodon Limnocyon Oxyaen odon Proviverra Sinopa Tritemnodon Machae roides

X

X

X

X

X

X

X

X X

Condylarthra

X X X

X

X X

Chiroptera

Du X X X

X

X

X X

X X

X

X

X

X X

X X X

X X

X

X X

X X X

(continued)

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APPENDIX A. (continued) Genus Br1a Br1b Bathyopsis Eobasileus Uintathe rium

Lagomorpha

Leporidae

Mytonolagus

Marsupialia

Didelphidae

Armintodelphys Copedelphys Peradectes H e r p e t o t h er i u m Nanodelphys

Mesonychia

Palaenodonta

Perissodactyla

Mesonychidae

Epoicotheriidae Metacheiromyidae Amynodontidae Brontotheriidae

Eomoropidae Equidae

Isectolophidae Family indet. Helaletidae

Hyracodontidae

Primates

Notharctidae

Omomyidae

X

X X X X

Tetrapassalus Brachianodon Metacheiromys

X X

Cantius Notharctus Smilodectes Ageitodendron Anaptomorphu s Ar ti m o n iu s Chipetaia Gazinius Hemiacodon Macrot arsius Mytonius Omomys Ourayia Shoshoniu s Sphacorhysis Trogolemur Uintanius Utahia Washakius Wyomom ys

Br3

Ui1a

Ui1b

Ui2

Ui3

Du

X

X

X X

X X X

X

X

X X X

X

X

X

X

X X

X X

X

X X X X

X X X

X X X

X X

X X

X

X X

X

Mesonyx Harpagolestes Hessoleste s

Amynodon Megalamynodo n Diplacodon Do licorhinus Duchesneodus Eotitanops M etarhinus Metatelmatheriu m Paleosyops Protitanotherium Mesatirhinus Sphenocoelus Sth en od ecte s Eomoropus Hyracotherium Orohippus Ep ih ip p u s Duchesnehippus Isectolophus Desmatotherium Colodon Helaletes Heptodon Dilophodon Epitriplopus Hyrachyus Hyracodon T r ip l op us

Br2

X X

X X X

X X X

X X X

X

X

X X

X

X

X

X

X X

X X

X

X

X

X

X

X

X X

X X

X

X

X

X X X

X

X X

X

X

X

X

X

X

X

X

X

X X

X

X

X X

X

X

X

X

X

X

X

X

X X

X X

X

X

X

X

X

X

X

X

X X

X X X

X X X

X X

X X

X

X

X

X X

X

X X

X X

X

X X

X

X

X

X

X

X

X

X X

X

X

X

X X X X X

X X X X

X X X X

X X

X

X

X

X (continued)

Middle Eocene rock units in the Bridger and Uinta Basins

Orde r

Family

APPENDIX A. (continued) Genu s Br1a Br1b

Plesiadapiformes

Microsyopsidae

Microsyops Ui n ta s o r e x

Rodentia

Allomyidae Eutypom yidae Family indent.

Spurimus Janimus Uintamys Natrona Acritoparamys Microparamys M y t on om y s Paramys Pseudotomu s Quadratomus Re ith r o p a r a m y s Thisbemys Uintaparam ys Knightomys Pauromys Sciuravus Taxymys Tillomys Mysops Pareumys M et a n oi a m y s Protadjidaumo Protopt yhcu s Griphomys Passaliscomys

Ischyromyidae

Sciuravidae

Cylindrodontidae E o m y i da e Protopychidae Family indent. Heliscomyidae Tillodontia

Esthonychidae

Taeniodonta Stylinodontidae Note: Modified from Gunnell et al. (2009).

Tillodon Trogosus Stylinodon

ACKNOWLEDGMENTS The authors wish to express their deepest gratitude to the many colleagues, students, and volunteers who have assisted with our field work in the Bridger and Uinta basins over the years. In particular, we are grateful for the participation of faculty, staff, and students from the University of Colorado Museum, Washington University in St. Louis, San Diego Natural History Museum, College of Charleston, and Lamar University. We also thank the Wyoming and Utah State Offices of the U.S. Department of Interior Bureau of Land Management, and the Kemmerer, Rock Springs and Vernal field offices. Without the support of the BLM, our field work would not have been possible. Finally, we thank Justin Strauss and Margaret Madsen for their assistance with the compilation of the field trip mileage log. REFERENCES CITED Alexander, J.P., and Burger, B.J., 2001, Stratigraphy and taphonomy of Grizzly Buttes, Bridger Formation, and the middle Eocene of Wyoming; Eocene biodiversity; unusual occurrences and rarely sampled habitats: Topics in Geobiology, v. 18, p. 165–196. Andersen, D.W., and Picard, M.D., 1972, Stratigraphy of the Duchesne River Formation (Eocene-Oligocene), northern Uinta Basin, Northwestern Utah: Bulletin of the Utah Geological and Mineral Survey, v. 97, p. 1–29.

X

X

163

Br2

Br3

Ui1a

Ui1b

X X

X X

X X

X

Ui2

Ui3

Du

X

X

X X X X

X X X

X

X X

X X X

X X X X X

X X

X X

X X X X

X X X X X X

X X X X X X

X

X X X X

X X X X

X X X

X

X

X X X

X

X

X X X X X X X

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X X X X

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X X

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X X

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X

X

X

Betts, C.W., 1871, The Yale College expedition of 1870: Harper’s New Monthly Magazine, v. 43, p. 663–671. Black, C.C., and Dawson, M.R., 1966, A review of late Eocene mammalian faunas from North America: American Journal of Science, v. 264, p. 321– 349, doi:10.2475/ajs.264.5.321. Bradley, W.H., 1931, Origin and microfossils of the Green River Formation of Colorado and Utah: U.S. Geological Survey Professional Paper 168, p. 58. Bradley, W.H., 1964, Geology of Green River Formation and associated Eocene rocks in southwestern Wyoming and adjacent parts of Colorado and Utah: U.S. Geological Survey Professional Paper 496-A, p. 1–86. Brand, L.R., 2007, Lacustrine deposition in the Bridger Formation: Lake Gosiute extended: The Mountain Geologist, v. 44, p. 69–77. Brand, L.R., Goodwin, H.T., Ambrose, P.D., and Buchheim, H.P., 2000, Taphonomy of turtles in the Middle Eocene Bridger Formation: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 162, p. 171–189, doi:10.1016/ S0031-0182(00)00111-5. Bryant, B., Naeser, C.W., Marvin, R.F., and Mehnert, H.H., 1989, Upper Cretaceous and Paleogene sedimentary rocks and isotopic ages of Paleogene tuffs, Uinta Basin, Utah: U.S. Geological Survey Bulletin 1787-J, p. 1–22 p. Cashion, W.B., 1974, Geologic map of the Southam Canyon quadrangle, Uintah County, Utah: U.S. Geological Survey Miscellaneous Field Studies Map, MF-579. Cashion, W.B., 1986, Geologic map of the Bonanza quadrangle, Uintah County, Utah: Miscellaneous Field Studies Map, MF-1865. Chadey, H.F., 1973, Historical Aspects of the Green River Basin, Wyoming, in 25th Field Conference on the geology and mineral resources of the greater Green River Basin, Wyoming: Wyoming Geological Association, p. 27–33. Cope, E.D., 1872, Descriptions of some new Vertebrata from the Bridger Group of the Eocene: Proceedings of the American Philosophical Society, v. 12, p. 460–465. Cope, E.D., 1873, On the extinct Vertebrata of the Eocene of Wyoming, observed by the expedition of 1872, with notes on the geology, in Hayden,

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The Geological Society of America Field Guide 21 2011

Middle Cryogenian (“Sturtian”) Pocatello Formation: Field relations on Oxford Mountain and the Portneuf area, southeast Idaho Joshua A. Keeley Paul K. Link Department of Geosciences, Idaho State University, Pocatello, Idaho 83209, USA

ABSTRACT Detailed mapping along the east face of Oxford Ridge in the southern Bannock Range, southeast Idaho determines the stratigraphic placement and lateral extent of strata in the Scout Mountain Member of the Neoproterozoic Pocatello Formation. The lower “transitional unit” overlies the Bannock Volcanic Member and consists of 70 m of massive diamictite with argillitic and vesicular basaltic clasts up to cobble size intercalated with thin metabasalt and hyaloclastite units. Overlying the transitional unit is a 150–190-m-thick, massive, brown-green to purple sandy diamictite with dominantly quartzose cobble clasts. Interbedded with this middle unit is a 60-m-thick epiclastic volcanic interval informally named the Oxford Mountain tuffite. An upper sandstone unit up to 100 m thick lies above the diamictite at the head of Fivemile Creek in the southern portion of the map area. The volcanic interval contains plagioclase-phyric volcanic lithic sandstone, porphyritic volcanic lithic fragments and rounded cobbles in tuffaceous diamictite and a reworked stratified lapilli-tuff. It is interstratified with quartzose and volcanogenic diamictite and can be traced along 5.5 km of strike. On Oxford Mountain, laser ablation–inductively coupled plasma mass spectrometry U-Pb zircon ages presented here and additional sensitive high-resolution ion microprobe ages constrain the underlying Bannock Volcanic Member to be 717–686 Ma and require that the overlying Scout Mountain Member is younger than 685 Ma.

INTRODUCTION Rift-related volcanic, diamictite, and carbonate-bearing strata of the Pocatello Formation record rifting of Rodinia during widespread regional glaciation (Ludlum, 1942; Crittenden et al., 1983; Link, 1983; Link et al., 1994; Smith et al., 1994; Lorentz et al., 2004; Corsetti et al., 2007). Similar successions

occur along the entire length of the North American Cordillera from the Death Valley region to the Yukon Territory and include local exposures of volcanic rocks overlain by 5–8 km of passive margin sediments (Fig. 1) (Link et al., 1993; Stewart and Poole, 1974). Constraining the precise timing and duration of rifting lies at the heart of paleogeographic and snowball earth models that seek to explain drastic tectonic and climatic fluctuations near the

Keeley, J.A., and Link, P.K., 2011, Middle Cryogenian (“Sturtian”) Pocatello Formation: Field relations on Oxford Mountain and the Portneuf area, southeast Idaho, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 167–182, doi:10.1130/2011.0021(07). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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Keeley and Link and the 709 ± 5 and 717 ± 4 Ma SHRIMP U-Pb ages from the Pocatello Formation (Fanning and Link, 2004). However, reanalysis of the former yielded a revised age of 686 ± 4 Ma (Fanning and Link, 2008), requiring further work that is currently in preparation by Keeley. A main focus of this field guide is to present new stratigraphic, structural and geochronologic relationships on Oxford Mountain of the southern Bannock Range (Figs. 2 and 3). Specifically, the results establish that within the Scout Mountain Member, volcanic sandstone interbedded with diamictite is <685 Ma. This member lies above the Bannock Volcanic Member, 112°30′ 43°00′ RIV SN

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onset of complex life (Moores, 1991; Hoffman, 1991; Hoffman et al., 1998; Karlstrom et al., 1999). Radiometric data and subsidence analyses along the North American Cordilleran miogeoclinal succession suggest a twostage Rodinian rift history with rifting after 720 Ma and thermal subsidence after 570 Ma (Colpron et al., 2002). Recent paleomagnetic revisions from Australian cratons allow a longer duration for the paleogeographic Southwest U.S.–East Antarctic (SWEAT) plate tectonic configuration (Moores, 1991), suggesting Rodinian rifting began after ca. 720 Ma; Colpron et al., 2002) with significant continental drift by 650 Ma (Li and Evans, 2011). Previous paleomagnetic data suggested continental breakup before 750 Ma (Wingate and Giddings, 2000). The Pocatello Formation is correlative to globally distributed middle Cryogenian “Sturtian” glacial rocks that demonstrate diachronous glaciation between the isotope dilution–thermal ionization mass spectrometry (ID-TIMS) age of 716.47 ± 0.24 Ma (Macdonald et al., 2010) and the Re-Os age of 643 ± 2.4 Ma (Kendall et al., 2006). Equivalent strata of the Edwardsburg Formation in central Idaho yield ages of ca. 685 Ma (Lund et al., 2003). This is consistent with the sensitive high-resolution ion microprobe (SHRIMP) age of 659.7 ± 5.3 Ma from a dropstonebearing tuff within the South Australian Sturtian glacial succession immediately below the Tapley Hill Formation (Fanning and Link, 2006). Hoffman and Li (2009) consider the chemical abrasion– thermal ionization mass spectrometry (CA-TIMS) U-Pb zircon age of 711.5 ± 0.3 Ma (Bowring et al., 2007) from the glacial Ghubrah Formation in Oman to be the most precise Sturtian age. The age is about 5 m.y. younger than the ID-TIMS age of 716.47 ± 0.24 Ma from Rapitan-correlative diamictites of the Mount Harper Group in the Yukon Territory (Macdonald et al., 2010)

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Figure 1. Map showing the location of diamictite-bearing successions along the Cordilleran belt and Pocatello Formation. From Fanning and Link (2004).

Figure 2. Location map showing the study area on Oxford Mountain in the southern Bannock Range, SE Idaho. From Link (1983). Stops 1 and 2 are located on Oxford Mountain; Stop 3 is located in Pocatello.

Oxford Mountain and the Portneuf area, southeast Idaho which is 717–686 Ma (Fanning and Link, 2004; 2008). Work in progress suggests that the span of zircon ages defined by laser ablation–inductively coupled plasma mass spectrometry (LAICPMS) (>1% precision) and SHRIMP (~1% precision) contains distinct zircon populations between 709 and 685 Ma (Keeley et al., 2010). Field mapping at the 1:12,000 scale was conducted over the summers of 2009 and 2010 along the length of Oxford Ridge. The study was mostly limited to the exposures of the Pocatello Formation along the east face in the Clifton, Weston Canyon and Oxford quadrangles (Fig. 4). Detailed mapping was designed to: (1) determine the stratigraphic relationships and lateral extent of the previously dated reworked lapilli-tuff from north of Clifton Basin (sample 06PL00 of Fanning and Link, 2004; 2008); (2) carefully map, describe and test the gradational contact between Bannock Volcanic and Scout Mountain Members of

General Neoproterozoic section, Pocatello area 4500

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Geologic Setting and Regional Stratigraphy of the Pocatello Formation The bedrock underlying southeast Idaho includes strata of the Pocatello Formation (Ludlum, 1942), the overlying Neoproterozoic-Cambrian Brigham Group (Link et al., 1987) (Figs. 1 and 3) overlain by Late Proterozoic to Ordovician rocks of the Sauk Sequence (Sloss, 1963). The Neoproterozoic Pocatello Formation is exposed in the Bannock and Pocatello ranges with minor outcrops in the southern Portneuf Range of southeast Idaho (Ludlum, 1942). The outcrops in Idaho are aligned N-S and extend ~120 km from

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Figure 3. Generalized stratigraphic sections. (A) Entire Neoproterozoic section in Pocatello area. (B) Pocatello Formation stratigraphy at Portneuf Narrows. (C) Pocatello Formation stratigraphy at Oxford Mountain. Modified from Fanning and Link (2004) based on new mapping at Oxford Mountain.

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Oxford Mountain and the Portneuf area, southeast Idaho the northern end of the Bannock Range at Moonlight Mountain to the southern end of the Bannock Range at Fivemile Creek, 15 mi south of Oxford Mountain (Fig. 2). The lower contact is never exposed. Part of the type section of the Pocatello Formation is located in an overturned limb of an east-vergent fold, south of China Peak and north of the Portneuf Narrows near Pocatello (Ludlum, 1942; Link, 1983). Here the 1500-m-thick Pocatello Formation is divided into the Bannock Volcanic Member (lower), the Scout Mountain Member (middle), and the upper member (upper) (Crittenden et al., 1971, 1983; Link, 1983) (Fig. 3). Lithologies include metabasalt (Bannock Volcanic Member), diamictitebearing siliciclastic rocks with minor carbonate (Scout Mountain Member), and phyllitic shale (upper member). In more detail, the Scout Mountain Member at Portneuf Narrows includes a lower diamictite (Stop 3), which is a locally stratified, green to brown, matrix-supported diamictite that contains clasts of mafic volcanic rocks, argillite and rare felsic volcanic rocks. Above this lies sandstone and conglomerate and an upper massive diamictite unit containing quartzitic, gneissic, granitic, and felsic volcanic clasts plus glacially striated quartzite clasts (Link, 1982a). Above the upper diamictite is a cap-carbonate unit: a pink, bedded dolomite and reworked dolomite chip breccia with interstratified sandstone and argillite. Well-exposed sections of this diamictite-cap carbonate sequence are also exposed above and north of Fort Hall Mine, south of Portneuf Gap, and provide a record of post-glacial depositional processes (see Dehler et al., this volume). A porcelaneous 10-cm-thick reworked green tuff (667 ± 5 Ma, SHRIMP concordia age; Fanning and Link, 2004) overlies the cap carbonate. The tuff lies below mediumbedded dark gray limestone interbedded with argillite that grades upwards into the argillitic upper member of the Pocatello Formation (Fanning and Link, 2004). Pocatello Formation on Oxford Mountain, Southern Bannock Range Due to exhumation associated with Cretaceous thrusting and Miocene-Pliocene extension, the study area along the east face of the Oxford Ridge exposes the deepest structural and lowest stratigraphic levels seen in the Idaho miogeocline (Fig. 4). Outcrops display at least 200 m of pillow basalt (Fig. 5A) and hyaloclastite (Fig. 5B) of the Bannock Volcanic Member and up to 250 m of overlying Scout Mountain Member diamictite (Figs. 5C–5E), with three stratigraphic units: a lower transitional unit of mainly diamictite, a middle extrabasinal diamictite (Fig. 5E) with interbedded plagioclase-rich volcanic lithic wacke, sand-

Figure 4. (A) Simplified geologic map of the Cache Valley, Malad Valley and Marsh Valley areas. New mapping of Oxford Ridge is bordered in black. Modified from Janecke et al., (2003). Inset after Long et al., (2006). (B) New geologic map of Oxford Ridge showing cross section transects and location of Stop 1 (see Fig. 8).

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stone and reworked lapilli-tuff (“the Oxford Mountain tuffite”; Fig. 5C), and an upper sandstone. The Oxford Mountain tuffite consists of up to 60 m (measured from cross sections) of volcanic sandstone, resedimented pyroclast-rich lapilli-tuff, and tuffaceous diamictite. Specific lithologies include non-stratified mafic volcaniclastic diamictite, non-stratified dacite-trachyte-bearing volcanic diamictite, stratified lapilli-tuff with soft sediment deformation, thin- to medium-bedded graded tuffaceous sandstone (67JK09, and 06PL00 of Fanning and Link, 2004; 2008) and crudely to nonstratified dacite-trachyte-bearing lapillistone. The tuffaceous strata are exposed along 5.5 km of strike above the low-angle Clifton Canyon fault. New mapping (Figs. 4, 6, and 8) demonstrates that the diamictite of the transitional unit, Scout Mountain Member on Oxford Ridge is in gradational contact above hyaloclastite of the Bannock Volcanic Member. Altered mafic volcanics and hyaloclastite of the Bannock Volcanic Member are punctuated by quartzose diamictite, and interbeds of hyaloclastite persist upsection until gradation into mafic clast and argillaceous diamictite of the Scout Mountain Member. Diversity of clasts from the Scout Mountain Member diamictite on Oxford Mountain suggests a mixture of intrabasinal and extrabasinal components. Bannock Range Structure The southern Bannock Range is located within the hanging wall of the Paris thrust, which may have slipped 60 km to the east relative to underlying strata (Armstrong and Oriel, 1965). Throughout the southern Bannock and Portneuf Ranges, Middle Cambrian strata are disconformably overlain by Miocene-Pliocene rocks of the Salt Lake Formation (Keller, 1963; Rodgers and Janecke, 1992). These relations indicate uplift and erosion of ~5000 m of strata as well as a subhorizontal orientation of Middle Cambrian and lower strata prior to the Miocene. To the west of the Bannock Range, Tertiary rocks overlie progressively younger Paleozoic rocks indicating a west-dipping ramp of the Paris thrust, the Malad Ramp (Rodgers and Janecke, 1992). Hanging wall transport up and over the ramp created an anticline known as the Cache-Pocatello culmination, which subsequently collapsed due to extension on the Bannock detachment system (Janecke et al., 2003; Carney and Janecke, 2005). The basal, middle Miocene part of the Salt Lake Formation is hypothesized to have been deposited in one supradetachment basin above the Bannock detachment system until Late Miocene (<10 Ma) breakup of the hanging wall into numerous normal fault-bounded basins throughout the region (Janecke and Evans, 1999; Janecke et al., 2003; Carney and Janecke, 2005). Many basin-bounding faults to the east and south of the map area are interpreted to be in the hanging wall of this laterally extensive, gently west-dipping Bannock detachment system which slipped in multiple phases between 10 and 4 Ma (Long et al., 2006). North-striking Miocene and Pliocene high-angle normal faults

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Figure 5. (A) Pillow basalt, Bannock Volcanic Member. Dashed white line marks contact between pillow basalt (below) and lobate flow (above). (B) Hyaloclastite and volcaniclastic rocks of the Bannock Volcanic Member. (C) Oxford Mountain reworked lapilli-tuff exhibiting slump folds. Dashed white lines mark planar and folded bedding planes. Pyroclast depresses laminae (arrow). (D) Scoured depositional contact (arrow) between extrabasinal conglomerate and diamictite on the Bannock Volcanic Member. (E) Stratified extrabasinal diamictite of the Scout Mountain Member. Arrow points toward the flow direction, marked by erosion on the leeward side of quartzite clasts.

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Figure 6. Cross sections A–D along Oxford Ridge from north (A) to south (D). (A) Near the northernmost extent of the Clifton fault with the Camelback Mountain Quartzite (CZu) in its hanging wall and the underlying and folded Clifton Canyon fault. (B) Near the northernmost extent of the Oxford Ridge anticline and subsequent folding of the New Canyon, Clifton and Clifton Canyon low-angle normal faults. (C) Back rotated, east-dipping Clifton fault cut by the structurally higher New Canyon fault. (D) WSW-dipping and back-tilted Clifton Canyon fault intruded by a possibly sheet-like Tertiary intrusion.

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are prevalent throughout the region and provide evidence of Basin and Range crustal extension. The southern Bannock Range is a Basin and Range horst bounded on the east by two imbricate east-dipping faults, the Dayton-Oxford fault and the East Dayton-Oxford fault, and bounded on the west by the west-dipping Deep Creek fault. In the footwall of the horst the older low-angle Clifton fault places Neoproterozoic-Cambrian quartzite of the Camelback Mountain Quartzite on the Neoproterozoic Pocatello Formation (Link, 1982a, 1982b; Carney et al., 2002; Carney and Janecke, 2005). The hanging wall of the Clifton fault is cut by a roughly parallel low-angle normal fault, the New Canyon fault, which places Ordovician to Cambrian carbonate strata on Brigham Group quartzite and the Pocatello Formation (Janecke and Evans, 1999). The New Canyon fault locally cuts out the Clifton fault and places Cambrian carbonate rocks directly on diamictite of the Pocatello Formation. Both low-angle faults strike NNW and dip ~15° to the WSW. Together they have accommodated ~15 km of extension (Carney, 2002; Carney and Janecke, 2005). Oxford Ridge Structure Subhorizontal faults on Oxford Ridge and the shallow east dips of foliation in footwall rocks, along with other flat-on-flat geometries reported to the south in the Weston Canyon quadrangle (Steely and Janecke, 2005; Steely et al., 2005), are part of the Bannock detachment system (Janecke and Evans, 1999; Janecke et al., 2003). The Bannock detachment system is hypothesized to have accommodated 50% extension (Carney and Janecke, 2005). Pocatello Formation strata in the footwall of the Bannock detachment on Oxford Mountain are strongly foliated, sheared and locally tightly folded beneath the Clifton fault. Foliation and bedding measurements along the ridge define the gentle Oxford Ridge anticline (Carney, 2002), whose fold axis is horizontal with a trend of 345° (Carney and Janecke, 2005). A topographic break in slope on the east face of Oxford Ridge occurs at the faulted contact between the underlying Bannock Volcanic Member and the overlying Scout Mountain Member. Bedding attitudes in the Scout Mountain Member above the fault are steeper than the east-dipping, roughly bedding-parallel, foliation in the Bannock Volcanic Member. Bedding-parallel foliation near the fault may be shear-related foliation that was rotated to low angles. A Tertiary mafic intrusion is intruded along the fault. Results of New Mapping New mapping has found at least two roughly 30-m-thick pillow basalt flows (Fig. 5A), separated by massive basalt flows and hyaloclastite (Fig. 5B) in the Bannock Volcanic Member exposed on the east face of Oxford Ridge. No primary rhyolite flows were found. One 2-m-thick clast-supported, quartzite cobble conglomerate appears interstratified within the lower portion of Bannock volcanics (65JK09). Stretched phyllite within the Bannock Volcanic Member, up to 10 m thick, exhib-

its alternating green-to-purple bands. The bands are 2 cm thick and foliation-parallel and are interpreted as fine grained turbidites deposited in a marine basin between basalt eruptions. A 1.5-m-thick tectonized purple marble above these strata may record primary carbonate deposition. The upper contact of the Bannock Volcanic Member is marked at the top of the last thick (>10 m) metabasalt or hyaloclastite. Clifton Canyon Road (Stop 1) shows the relations of interstratified white-pebble diamictite of the basal Scout Mountain Member and volcaniclastic rocks and hyaloclastite of the Bannock Volcanic Member. Elsewhere along Oxford Ridge, the contact between Bannock Volcanic and Scout Mountain Members is faulted by the structurally highest fault in the footwall of the Clifton fault, here named the Clifton Canyon fault after its southernmost exposure east of Clifton Basin (Stop 1). Just below the Clifton Canyon fault north of Davis Basin in section 17 of the Clifton quadrangle, extrabasinal diamictite and conglomerate of the Scout Mountain Member lie above a scoured but depositional contact on volcaniclastic rocks of the Bannock Volcanic Member (agglomerate facies of Link, 1982a) (Fig. 5D). Three mappable and intertonguing stratigraphic units have been identified in the Scout Mountain Member on Oxford Mountain (Fig. 3). The lower “transitional unit” Link (1982a), consists of 70 m of massive diamictite with argillitic and vesicular basaltic clasts up to cobble size. Trace element geochemistry of these clasts is similar to the Bannock Volcanic Member, suggesting their proximal derivation. Mafic volcanic clasts make up 80% of the clast composition (Link, 1982a; this study). The remaining clasts are dominantly quartzite with sparse basement clasts. This unit and its lower contact are best preserved along Clifton Road (Stop 1) where hyaloclastite is intercalated with a white-pebble diamictite and intraclastic diamictite. Overlying the transitional unit is a 150–190-m-thick, massive, brown-green to purple, sandy, locally volcaniclastic diamictite with dominantly quartzose cobble clasts (Fig. 5E). Interbedded within this diamictite is the Oxford Mountain tuffite, which contains up to 60 m of proximal to medial subaqueous volcaniclastic rocks. Other clast lithologies in the diamictite include porphyritic volcanic rocks, chloritic argillite chips, basement gneiss and basalt. This heterolithologic unit makes up the dominant lithology on Oxford Ridge and is thickest (200 m thick) at Fivemile Canyon. Locally within this diamictite is 25 m of medium- to thick-bedded trough-cross bedded sandstone. Similar sandstone up to 100 m thick forms an upper unit that lies above the diamictite at the head of Fivemile Creek. Facies within the Oxford Mountain tuffite include, generally bottom to top: (1) mafic volcanic diamictite, comprising pebble to cobble, angular to subrounded vesicular and non-vesicular basalt; (2) non-stratified dacite-trachyte-bearing volcanic diamictite including rounded to well-rounded pebbles, cobbles and rare boulders of gray to purple porphyritic dacite-trachyte epiclasts; (3) stratified and reworked lapilli-tuff with soft sediment deformation; (4) thin- to medium-bedded, graded tuffaceous sandstone which is generally between non-stratified volcanic diamictite

Oxford Mountain and the Portneuf area, southeast Idaho units (67JK09); (5) crudely to non-stratified dacite-trachyte– bearing volcanic diamictite including angular to subrounded, white to dark gray aphanitic clasts that commonly show light and dark rinds. Dark rinds are metamorphic biotite rims whereas light rinds are interpreted to be the result of carbonate and chloritic alteration. A sixth facies includes feldspathic sandstone and intraformational conglomerate and breccia wedges near the top of the unit. Interbedded mafic volcanic diamictite and tuffaceous units are interpreted to record reworking of the Bannock Volcanic Member during tuffite deposition. Felsic cobbles and poorly sorted reworked lapilli-tuff suggest a proximal derivation of felsic volcanic clasts. The tuffite is not exposed in the southern portion of the map area in Fivemile Canyon. However, sample 75JK09 (Figs. 4 and 7) taken from the lowest exposed diamictite there has a prominent young zircon population at ca. 685 Ma. This may imply that the Oxford Mountain tuffite occurs near this interval yet has been reworked beyond recognition by diamictiteforming debris flow processes. Oxford Ridge does not expose Neoproterozoic carbonate or dark phyllitic strata seen in the upper Scout Mountain Member at Portneuf Narrows. If correlative strata existed, they have been omitted along the Clifton fault (Figs. 4 and 6). Diamictite correlation between Portneuf Narrows and Oxford Ridge is still poorly constrained. However, the Oxford Mountain tuffite can be loosely correlated to the Scout Mountain Member below the cap carbonate due to the lack of carbonate and phyllite on Oxford Ridge. New LA-ICPMS Results U-Pb LA-ICPMS isotope data was collected from a total of 464 detrital zircon grains from 8 samples along Oxford Ridge (Figs. 4 and 7). Samples were analyzed at the University of Arizona Laserchron lab following analytical methods outlined in Gehrels et al. (2008) and references therein. Precision on these analyses is >1% or a standard deviation of ±7 m.y. in a 700 Ma sample. Analyses less than 80% concordant were eliminated. More precise analyses by SHRIMP (1% precision) and TIMS (0.1% precision) are in progress to isolate the grain-population shown in Figure 7F. Three samples are from the southernmost portion of the map area at Fivemile Canyon. These include a lowermost transitional diamictite (75JK09; Fig. 7B), a quartzite clast in that diamictite (74JK09; Fig. 7C) and a higher diamictite sample just below the Clifton fault (73JK09; Fig. 7D). Both diamictites contain the same age-peaks with the largest zircon-age population at 1.7 Ga, a small population at 2.48 Ga and a low-amplitude spread between 1.0 and 1.6 Ga. 75JK09 has a significant population at ca. 685 Ma that diminishes upsection, which provides a maximum age constraint for diamictite deposition. The quartzite clast (74JK09, Fig. 7) has zircons all older than 1500 Ma, with a major population at 1.7 Ga, duplicating the zircon distribution in the diamictite within which it is found.

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Samples from the east face of Oxford Ridge include a quartzite cobble and sandstone from a cobble conglomerate in the Bannock Volcanic Member, the stratigraphically lowest sample of the set (65JK09; Fig. 7A), one granite clast from diamictite in scoured contact with the underlying Banock Volcanic Member (24JK09; Fig. 7A inset), one sample of volcanic sandstone (68JK09; Fig. 7E) and one volcanic diamictite (67JK09; Fig. 7F) within the middle diamictite unit. The conglomerate in the Bannock Volcanic Member (65JK09; Fig. 7A) has a ca. 685 Ma zircon population that provides a maximum age for most of the Bannock Volcanic Member on Oxford Mountain. That this is a reasonable estimate of the age of the Bannock Volcanic Member is supported by a SHRIMP age of 689 ± 4 Ma from a porphyritic rhyolite cobble from the cobble conglomerate submember at Portneuf Narrows (Keeley et al., 2010). The other zircons in this lowest extrabasinal siliciclastic interval are age equivalent to, and probably sourced from the Archean Wyoming craton (2.4– 3.2 Ga) and the 1.7 Ga Mojave-Yavapai-Mazatzal province. The granite clast (24JK09; inset Fig. 7A), removed from the Scout Mountain Member directly above the scoured contact above the Bannock Volcanic Member, yields ages >2.5 Ga, evidence of derivation from the southern edge of the Archean Wyoming Craton. Samples 73JK09, 74JK09, 75JK09 and 68JK09 all pass the Kolmogorov-Smirnoff (K-S) test using the cumulative distribution function while taking the standard error of each analysis into consideration (Press et al., 1986). The successful K-S test thus means all four samples are statistically indistinguishable and supports a similar provenance. SHRIMP and TIMS geochronology on the ca. 685 Ma zircons is in progress. The Oxford Mountain plagioclase-rich volcanic diamictite (67JK09; Fig. 7F) contains three or more zircon sub-populations (ca. 685, ca. 695, and 702 ± 5 Ma) which overlap with SHRIMP ages for 705 Ma grain populations from the lower diamictite of Portneuf Narrows (Keeley et al., 2010). The volcanic sandstone (68JK09; Fig. 7E) exhibits the 685 Ma peak mixed with Laurentian provenance including 1.1–1.3 Ga (Grenville), 1.3–1.5 Ga (Mid-continent Granite) and 1.6–1.9 Ga (MojaveYavapai-Mazatzal province) age components. The inset in figure 7E shows a quartzite clast (3JK09) with exclusive 1.7 Ga provenance. Zircons of mixed Laurentian provenance could also be sourced from reworking of the early Cryogenian Uinta Mountain Group and Big Cottonwood Formation to the southeast (Dehler et al., 2010). The lack of abundant Archean grains in the volcanic sandstone (Fig. 7E) suggests a provenance shift to dominantly Paleoproterozoic sources. Based on the similar mixed Laurentian provenance seen in detrital zircon plots between the Fivemile Canyon area and Oxford Ridge, the diamictite along strike appears to have been deposited in one basin. ROAD LOG The road log begins on W. Oneida Street (Idaho Rt. 36) in Preston, ~26 mi north of Logan and 65 mi south of Pocatello.

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Oxford Mountain and the Portneuf area, southeast Idaho Incremental and cumulative mileage is given from the intersection with Main Street. Stop 1 (Fair Weather Stop Best Made after May; Four-Wheel-Drive Vehicle is Recommended) Mileage Incr. Cum. 0.0

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Drive west on W. Oneida Street (Highway 36). To the southwest is the low ridge, Rattlesnake Ridge, of the southern end of the Bannock Range in the Weston Canyon quadrangle. Bounding the ridge on the east is the Pliocene-Quaternary(?) Dayton-Oxford fault. The flat ridge crest marks the location and

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rough orientation of the west-dipping lowangle Clifton fault. To the northwest, Weston and Buck Peaks are visible in the Clifton quadrangle and further to the north stands Oxford Peak at 9265 ft. Highway 36 crosses the Bear River. Leave Highway 36 by bearing right in the town of Dayton on County Road DI (a.k.a. Oxford Highway). Enter the town of Clifton and take left onto W. 1st Street. Stay left onto gravel Cemetery Road. Road turns into Clifton Creek Road. The road is now in the footwall of the East DaytonOxford fault and in the hanging wall of the Dayton-Oxford fault. Beside the road to the

Figure 8. New map of Clifton Basin. Clifton Basin is bounded on the west by the east-dipping, low-angle Clifton Basin fault (CBF) and on the east by the WSW-dipping, low-angle Clifton fault (CF). The Clifton Canyon fault (CCF) is in the footwall of the Clifton fault. The range bounding fault, the Oxford-Dayton fault and the lowangle Clifton fault cuts Tertiary strata of the Salt Lake Formation. See Figure 4B for figure location. CZcm— Neoproterozoic Camelback Mountain Formation (Brigham Group); Clb—Cambrian Lead Bell Formation; Cbl— Cambrian Blacksmith Formation. See Figure 10 for detailed stratigraphy of the Pocatello Formation.

Figure 9. Sketch of transitional unit outcrops along Clifton Basin Road (Stop 1) displaying intertonguing diamictite and hyaloclastite of the lower transitional unit. Foliation strikes and dips are noted at each outcrop and are broadly folded about an axis of 309° plunging 26° to the south.

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north are east-dipping Tertiary beds of the Salt Lake Formation. South of the road are Neoproterozoic-Cambrian strata of the Camelback Mountain Quartzite. 0.8 13.7 The road crosses Clifton Creek. Stay right on public road. 0.8 14.5 The road crosses Clifton Creek. The road is now in the footwall of the DaytonOxford fault and begins to switchback up Oxford Ridge. 0.4 14.9 Pass Davis Lodge on right. Stay left on public road. 0.9 15.8 Crest hill and open the gate to pass. Please close the gate behind you. See Figure 8 for a map of the area. Outcrops to the left are shown in Figure 9. 0.5 16.3 Stop 1. Enter Clifton Basin by crossing the Clifton fault and park near the corral (Fig. 8). The purpose of this stop is to demonstrate the gradational contact between the Bannock Volcanic and Scout Mountain members. The view north from Clifton Basin displays Neoproterozoic-Cambrian rocks of the Brigham Group, Cambrian carbonate rocks, and Tertiary strata. The closest ridge to the northeast shows the Clifton fault placing Tertiary strata on the diamictite of the Pocatello Formation. To the south, topography is flat and west-dipping, marking the planar, low-angle Clifton fault. Begin hike by walking back to the gate on Clifton Basin Road. Total hike is 0.5 mi. Beyond and to the south of the gate is a weathered contact between the underlying metabasalt of the Bannock Volcanic member and overlying quartzite-clast–bearing Scout Mountain Member, where rocks on both sides are heavily foliated. This is the southern extent of the Clifton Canyon fault at its tip, where there is little to no offset. One can distinguish between volcaniclastic rocks bearing no detrital quartz component and the brown quartzose diamictite near the gate. The outcrop at the gate is a sandy, white-quartzite granule-pebble diamictite (Fig. 10). The next outcrop along the road displays a high-angle, lowoffset, down-to-the-west fault that is traceable across Clifton Canyon to the north. The fault is entirely within hyaloclastite; offset is at least 25 m based on an offset quartz vein to the north. Walk westward along the road distinguishing between outcrops of mafic-volcaniclastic pebble conglomerate, hyaloclastite and quartzose diamictite. These outcrops demonstrate the gradational contact between the Bannock Volcanic and Scout Mountain Members. Near Clifton Basin, hyaloclastite units grade upward into intrabasinal mafic-volcaniclastic diamictite and then into extrabasinal diamictite. Based on the occurrence of interbedded volcaniclastic, hyaloclastite, and quartzose units, the rocks were mapped as the transitional unit (lower Scout Mountain Member). Walk back toward Clifton Basin, cross the small creek near the corral and head northeast to climb the ridge to the slope break (6560 ft elevation). Walk along the slope break to the northeast,

black medium- to finegrained sandstone green clast-poor fragmental phyllitic diamictite coarse-grained volcaniclastics foliated brown mafic-chip diamictite greenschist phyllite Metabasalt 195, 17 280, 13 210, 44 233, 33 243, 25

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Oxford Mountain and the Portneuf area, southeast Idaho then north and notice the subtle change in lithology from quartzite-rich diamictite above the slope break and transitional, volcanic-clast-rich diamictite below. Rocks at this location are similar to those at field trip Stop 3, near Pocatello. 4.5 20.8 Return to vehicles. Drive back down Clifton Creek Road to the town of Clifton and take a left on County Road D1 (Oxford Highway). Head north observing cliffs to the west composed of Bannock Volcanic Member pillow basalts. 5.7 26.5 Enter the town of Oxford. 0.2 26.7 Take a left on Oxford Loop Road (dirt). 2.6 29.3 Stop 2. Park along the road near the intersection of Oxford Creek Road. This roadside stop displays a broad glimpse of the stratigraphy and structure of the Oxford Peak area, to provide a con-

text for the newly dated Oxford Mountain tuffite. The view to the W and WSW shows the cliffs above Oxford Basin (Fig. 11). The view to the SW and SSW shows the flat-topped eastward prong of Oxford Ridge. The flat top is made up of OrdovicianCambrian strata in the hanging wall of the New Canyon fault, which soles into the Clifton fault at this location. A topographic break below the New Canyon fault marks the location of the Clifton Canyon fault, above which lies the Oxford Mountain tuffite. Several more lineaments along the ridge in the footwall mark the position of moderate- to low-angle faults bounding slide blocks that are interpreted to have slipped during or after the PlioceneQuaternary(?) phase of Long et al. (2006). The Clifton fault is exposed on the high cliffs of Oxford Peak between the white quartzite of the Camelback Mountain Quartzite and the dark green metabasalt of the Bannock Volcanic Member of the Pocatello Formation. Footwall strata

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Figure 11. (A) Annotated sketch of Oxford Ridge and Oxford Peak from the junction of Oxford Loop Road and Oxford Creek Road. Note locations of low-angle, progressively higher, Clifton Canyon and Clifton faults, as well as high-angle Basin and Range faults, the Oxford Basin fault and range-bounding Dayton-Oxford fault. Zpb—Pocatello Formation, Bannock Volcanic Member; Zps—Pocatello Formation, Scout Mountain Member; Zpso—Pocatello Formation, Scout Mountain Member, Oxford Mountain tuffite; CZcm—Camelback Mountain Quartzite; OCu— Ordovician to Cambrian undifferentiated; Qls—Quaternary landslide; ORA—Oxford Ridge anticline. Refer to cross sections in Figure 6. (B) Photograph of Oxford Peak and Oxford Basin from junction of Oxford Loop Road and Oxford Creek Road. Same view as in A.

correlative to these basalts have been downdropped by the Oxford Basin fault (Figs. 4 and 6), a Basin-and-Range fault synthetic to the Dayton-Oxford fault. The low-lying hills also contain flat tops that are often capped by extensional klippen of the Clifton fault. The northernmost exposure of the Oxford Mountain tuffite and volcanic sandstone is visible between the Clifton and the Clifton Canyon faults at the southern end of the cliffs (Fig. 11). 2.1 31.4 Return to vehicles. Continue the field trip by taking a right onto Oxford Creek Road and driving to County Road D1 (Oxford Highway). 2.7 33.4 Stay left on Oxford Highway. 4.3 37.7 Take a left onto Highway 91 north to Pocatello. 12.2 49.9 Take a right (north) onto Interstate Highway 15. 21.0 70.9 Pass through the town of Inkom. The Pocatello Range on both sides of the road is made up of east-dipping NeoproterozoicCambrian rocks of the Brigham Group and underlying Blackrock Limestone. 6.2 77.1 Exit the freeway using the Portneuf Area exit 63. Turn right at the end of the off-ramp. The road is now in the hanging wall of the westdipping Fort Hall Mine fault. To the east,

across the fault, are east-tilted Pocatello Formation, Scout Mountain Member rocks in the footwall of the Fort Hall Mine fault. 1.5 78.6 Take a right onto N. Fort Hall Mine Road and go straight at the first stop sign. 0.5 79.1 Stop 3. Turn left onto a dirt road immediately before the Fort Hall Mine Landfill and park at the first turnout before the road turns right. The purpose of this stop is to compare the lower diamictite here to the Clifton Creek Road diamictite at Stop 1. Climb above Pocatello Formation–derived Tertiary strata to an elevation between 4760 ft and 5500 ft and observe the lower diamictite. Two samples 62JK09 (705 ± 5 Ma) and 63JK09 (704 ± 5 Ma) from this interval (Fig. 3B) have zircon populations with SHIRMP concordia ages consistent with age-populations in the Oxford Mountain volcanic sandstone (Keeley et al., 2010). See Dehler et al. (this volume) for nearby stops farther south exploring the depositional settings of the post-glacial dolomitic cap carbonate. CONCLUSIONS Detailed mapping along the east face of Oxford Ridge has constrained the stratigraphic placement and lateral extent of previously dated strata in the Scout Mountain Member of the

Oxford Mountain and the Portneuf area, southeast Idaho Neoproterozoic Pocatello Formation (Fanning and Link, 2004; 2008). The results support a gradational contact between the underlying Bannock Volcanic Member and the overlying Scout Mountain Member (Link, 1982a). Indistinguishable detrital zircon U-Pb age spectra between Fivemile Canyon and Oxford Ridge suggest a similar mixed Laurentian source and may suggest deposition within a single, newly-formed rift basin. U-Pb analysis has also yielded a young zircon population of ca. 685 Ma. More precise SHRIMP and TIMS geochronology on this population is in progress. ACKNOWLEDGMENTS This research was supported by National Science Foundation grant EAR-0819759. We are grateful for reviews by David Rodgers, Carol Dehler, and James Evans. REFERENCES CITED Armstrong, F.C., and Oriel, S.S., 1965, Tectonic development of the IdahoWyoming thrust belt: The American Association of Petroleum Geologists Bulletin, v. 49, p. 1847–1866. Bowring, S.A., Grotzinger, J.P., Condon, D.J., Ramezani, J., Newall, M.J., and Allen, P.A., 2007, Geochronologic constraints on the chronostratigraphic framework of the Neoproterozoic Huqf Supergroup, Sultanate of Oman: American Journal of Science, v. 307, no. 10, p. 1097–1145, doi:10.2475/10.2007.01. Carney, S.M., 2002, Evolution of a Miocene-Pliocene low angle normal-fault system in the southern Bannock Range, southeast Idaho [M.S. thesis]: Logan, Utah State University, 177 p. Carney, S.M., and Janecke, S.U., 2005, Excision and the original low dip of the Miocene-Pliocene Bannock detachment system, SE Idaho: Northern cousin of the Sevier Desert detachment?: Geological Society of America Bulletin, v. 117, p. 334–353, doi:10.1130/B25428.1. Carney, S. M., Janecke, S. U., Oriel, S. S., Evans, J. C., and Link, P. K., 2002, Geologic map of the Clifton quadrangle, Franklin and Oneida Counties, Idaho, Idaho Geological Survey Technical Report T-03-4, scale 1:24,000. 2 plates. Colpron, M., Logan, J., and Mortensen, J., 2002, U-Pb zircon age constraint for late Neoproterozoic rifting and initiation of the lower Paleozoic passive margin of western Laurentia: Canadian Journal of Earth Sciences, v. 39, no. 2, p. 133–143, doi:10.1139/e01-069. Corsetti, F., Link, P.K., and Lorentz, N.J., 2007, δ13C Chemostratigraphy of the Neoproterozoic succession near Pocatello, Idaho, U.S.A.: Implications for glacial chronology and regional correlations, in Link, P.K., and Lewis, R.S., eds., Proterozoic geology of western North America and Siberia: SEPM Special Publication 86, p. 193–205. Crittenden, M., Schaeffer, F., Trimble, D., and Woodward, L., 1971, Nomenclature and correlation of some upper Precambrian and basal Cambrian sequences in western Utah and southeastern Idaho: Geological Society of America Bulletin, v. 82, no. 3, p. 581–601, doi:10.1130/0016 -7606(1971)82[581:NACOSU]2.0.CO;2. Crittenden, M.D., Jr., Christie-Blick, N., and Link, P., 1983, Evidence for two pulses of glaciation during the late Proterozoic in northern Utah and southeastern Idaho: Geological Society of America Bulletin, v. 94, no. 4, p. 437–450, doi:10.1130/0016-7606(1983)94<437:EFTPOG>2.0.CO;2. Dehler, C., Fanning, C., Link, P., Kingsbury, E., and Rybczynski, D., 2010, Maximum depositional age and provenance of the Uinta Mountain Group and Big Cottonwood Formation, northern Utah: Paleogeography of rifting western Laurentia: Geological Society of America Bulletin, v. 122, no. 9–10, p. 1686, doi:10.1130/B30094.1. Dehler, C., et al., 2011, this volume, New descriptions of the cap dolostone and associated strata, Neoproterozoic Pocatello Formation, southeastern Idaho, USA, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes

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of the U.S. Cordillera: Geological Society of America Field Guide 21, doi:10.1130/2011.0021(08). Fanning, C.M., and Link, P., 2004, U-Pb SHRIMP ages of Neoproterozoic (Sturtian) glaciogenic Pocatello Formation, southeastern Idaho: Geology, v. 32, no. 10, p. 881, doi:10.1130/G20609.1. Fanning, C.M., and Link, P.K., 2006, Constraints on the timing of the Sturtian glaciation from southern Australia: i.e., for the true Sturtian: Geological Society of America Abstracts with Programs, v. 38, no. 7, p. 115. Fanning, C.M., and Link, P.K., 2008, Age constraints for the Sturtian Glaciation; data from the Adelaide Geosyncline, South Australia and Pocatello Formation, Idaho, USA, in Gallagher, S.J., and Wallace, M.W., eds., Neoproterozoic extreme climates and the origin of early metazoan life: Geological Society of Australia Extended Abstracts No. 91, p. 57–62. Gehrels, G., Valencia, V., and Ruiz, J., 2008, Enhanced precision, accuracy, efficiency, and spatial resolution of U-Pb ages by laser ablation–multicollector– inductively coupled plasma–mass spectrometry: Geochemistry Geophysics Geosystems, v. 9, no. 3, p. Q03017, doi:10.1029/2007GC001805. Hoffman, P., 1991, Did the Breakout of Laurentia Turn Gondwanaland InsideOut?: Science, v. 252, no. 5011, p. 1409–1412, doi:10.1126/science .252.5011.1409. Hoffman, P., and Li, Z., 2009, A palaeogeographic context for Neoproterozoic glaciation: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 277, no. 3-4, p. 158–172, doi:10.1016/j.palaeo.2009.03.013. Hoffman, P., Kaufman, A., Halverson, G., and Schrag, D., 1998, A Neoproterozoic Snowball Earth: Science, v. 281, no. 5381, p. 1342, doi:10.1126/ science.281.5381.1342. Janecke, S.U., and Evans, J.C., 1999, Folded and faulted Salt Lake Formation above the Miocene to Pliocene New Canyon and Clifton detachment faults, Malad and Bannock Ranges, Idaho: Field trip guide to the Deep Creek half graben and environs, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the Geology of Eastern Idaho: Pocatello, Idaho Museum of Natural History, p. 71–96. Janecke, S.U., Carney, S., Perkins, M., Evans, J., Link, P.K., Oaks, R.Q., Jr., and Nash, B., 2003, Late Miocene–Pliocene detachment faulting and superimposed Pliocene-Pleistocene Basin and Range extension inferred from dismembered rift basins filled with the Salt Lake Formation, S.E. Idaho, in Raynolds, R.G., and Flores, R.M., eds., Cenozoic Systems of the Rocky Mountain Region: Rocky Mountain Section, SEPM, p. 369–406. Karlstrom, K.E., Harlan, S.S., Williams, M.L., McLelland, J., Geissman, J.W., and Åhäll, K.-I., 1999, Refining Rodinia: Geologic evidence for the Australian– western U.S. connection in the Proterozoic: GSA Today, v. 9, no. 5, p. 1–7. Keeley, J.A., Link, P.K., and Fanning, C.M., 2010, Structure, stratigraphy and geochronology of the Bannock Volcanic and Scout Mountain Members, Pocatello Formation, SE Idaho: Gradational contact constrains timing of Rodinian rifting and Sturtian glaciation: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 307. Keller, A.S., 1963, Structure and stratigraphy behind the Bannock thrust in parts of the Preston and Montpelier quadrangles, Idaho (Ph.D. thesis): New York, Columbia University, 205 p. Kendall, B.S., Creaser, R.A., and Selby, D., 2006, Re-Os geochronology of postglacial black shales in Australia: constraints on the timing of “Sturtian” glaciation: Geology, v. 34, p. 729–732, doi:10.1130/G22775.1. Li, Z., and Evans, D., 2011, Late Neoproterozoic 40° intraplate rotation within Australia allows for a tighter-fitting and longer-lasting Rodinia: Geology, v. 39, no. 1, p. 39, doi:10.1130/G31461.1. Link, P.K., 1982a, Geology of the upper Proterozoic Pocatello Formation, Bannock Range, southeastern Idaho [Ph.D. thesis.]: Santa Barbara, University of California, Santa Barbara, 131 p. Link, P.K., 1982b, Structural Geology of the Oxford Peak and Malad Summit quadrangles, Bannock Range, southeastern Idaho, in Powers, R.B., ed., Geologic Studies of the Cordilleran thrust belt: Denver, Colorado, Rocky Mountain Association of Geologists, p. 851–858. Link, P.K., 1983, Glacial and tectonically influenced sedimentation in the Upper Proterozoic Pocatello Formation, southeastern Idaho, in Miller, D.M., Todd, V.R., and Howard, K.A., eds., Tectonic and Stratigraphic Studies in the Eastern Great Basin: Geological Society of America Memoir 157, p. 165–181. Link, P.K., Jansen, S.T., Halimdihardja, P., Lande, A.C., and Zahn, P.D., 1987, Stratigraphy of the Brigham Group (Late Proterozoic–Cambrian), Bannock, Portneuf, and Bear River Ranges, southeastern Idaho, in Miller, W.R., ed., The Thrust Belt Revisited, 38th Annual Wyoming Geological Association Guidebook. p. 133–148.

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Link, P.K., Christie-Blick, N., Stewart, J.H., Miller, J.M.G., Devlin, W.J., and Levy, M.E., 1993, Late Proterozoic strata of the United States cordillera, in Reed., J., Simms, P., Houston, R., Rankin, D., Link, P., Van Schmus, R., and Bickford, P., eds., Precambrian of the conterminous United States: Boulder, Colorado, Geological Society of America, The Geology of North America, v. C-3, p. 536–558. Link, P.K., Miller, J.M.G., and Christie-Blick, N., 1994, Glacial-marine facies in a continental rift environment: Neoproterozoic rocks of the western United States Cordillera, in Deynoux, M., Miller, J.M.G., Domack, E.W., Eyles, N., Fairchild, I.J., and Young, G.M., eds., International Geological Correlation Project 260: Earth’s Glacial Record, Cambridge, U.K., Cambridge University Press, p. 29–59. Long, S.P., Link, P.K., Janecke, S.U., Perkins, M.E., and Fanning, C.M., 2006, Multiple phases of Tertiary extension and synextensional deposition of the Miocene-Pliocene Salt Lake Formation in an evolving supradetachment basin, Malad Range, southeast Idaho, U.S.A: Rocky Mountain Geology, v. 41, p. 1–27, doi:10.2113/gsrocky.41.1.1. Ludlum, J., 1942, Pre-Cambrian formations at Pocatello, Idaho: The Journal of Geology, v. 50, p. 85–95, doi:10.1086/625027. Lund, K., Aleinikoff, J., Evans, K., and Fanning, C., 2003, SHRIMP U-Pb geochronology of Neoproterozoic Windermere Supergroup, central Idaho: Implications for rifting of western Laurentia and synchroneity of Sturtian glacial deposits: Geological Society of America Bulletin, v. 115, no. 3, p. 349–372, doi:10.1130/0016-7606(2003)115<0349:SUPGON>2.0.CO;2. Macdonald, F., Schmitz, M., Crowley, J., Roots, C., Jones, D., Maloof, A., Strauss, J., Cohen, P., Johnston, D., and Schrag, D., 2010, Calibrating the Cryogenian: Science, v. 327, no. 5970, p. 1241, doi:10.1126/ science.1183325. Moores, E., 1991, Southwest U.S.–East Antarctic (SWEAT) connection; a hypothesis: Geology, v. 19, no. 5, p. 425–428, doi:10.1130/0091 -7613(1991)019<0425:SUSEAS>2.3.CO;2. Press, W.H., Flannery, B.P., Tenkolsky, S.A., and Vetterling, W.T., 1986, Numerical recipes: Cambridge, Cambridge University Press, p. 818.

Rodgers, D.W., and Janecke, S.U., 1992, Tertiary paleogeographic maps of the western Idaho Wyoming-Montana thrust belt, in Link, P.K., Kuntz, M., and Platt, L.B. eds., Regional Geology of Eastern Idaho & Western Wyoming: Geological Society of America Memoir 179, p. 83–94. Sloss, L., 1963, Sequences in the cratonic interior of North America: Geological Society of America Bulletin, v. 74, no. 2, p. 93, doi:10.1130/0016 -7606(1963)74[93:SITCIO]2.0.CO;2. Smith, L.H., Kaufman, A.J., Knoll, A.H., and Link, P.K., 1994, Chemostratigraphy of predominantly siliciclastic Neoproterozoic successions: A case study of the Pocatello Formation and lower Brigham Group, Idaho, USA: Geological Magazine, v. 131, p. 301–314, doi:10.1017/ S0016756800011079. Steely, A.N., and Janecke, S.U., 2005, Geologic map of the Weston Canyon 7.5′quadrangle, Oneida and Franklin Counties, Idaho: Technical Report T-05-3: Idaho Geological Survey, scale 1:24,000. Steely, A.N., Janecke, S.U., Long, S.P., Carney, S.C., Oaks, R.Q., Jr., Langenheim, V.E., and Link, P.K., 2005, Evolution of a Late Cenozoic supradetachment basin above a flat-on-flat detachment with a folded lateral ramp, SE Idaho, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 169–198. Stewart, J.H., and Poole, F.G., 1974, Lower Paleozoic and uppermost Precambrian strata of the Cordilleran Miogeocline, Great Basin, Western United States, in Dickinson, W.R., ed., Tectonics and Sedimentation: SEPM Special Publication 1, p. 28–57. Wingate, M.T.D., and Giddings, J.W., 2000, Age and palaeomagnetism of the Mundine Well dyke swarm, Western Australia: Implications for an Australia-Laurentia connection at 755 Ma: Precambrian Research, v. 100, p. 335–357, doi:10.1016/S0301-9268(99)00080-7.

MANUSCRIPT ACCEPTED BY THE SOCIETY 20 MARCH 2011

Printed in the USA

The Geological Society of America Field Guide 21 2011

New descriptions of the cap dolostone and associated strata, Neoproterozoic Pocatello Formation, southeastern Idaho, USA Carol M. Dehler Utah State University, Department of Geology, 4505 Old Main Hill, Logan, Utah 84322-4505, USA Kathleen Anderson Department of Geological Sciences, Brigham Young University, S389 ESC Provo, Utah 84602, USA Robin Nagy Utah State University, Department of Geology, 4505 Old Main Hill, Logan, Utah 84322-4505, USA

ABSTRACT An ~90-m-thick interval of mixed siliciclastic-carbonate strata, including a cap dolostone unit, overlies diamictite of the upper Scout Mountain Member of the Pocatello Formation in the Fort Hall Mine area south of Portneuf Narrows, southeastern Idaho, and is ≤ ca. 665 Ma. Six facies comprise this interval: silty sandstone (reworked diamictite matrix), laminated dolomite, dolomite and sandstone, sandstone, dolomite-chip breccia, and argillite and limestone. Sedimentary structures and bedding geometries of facies indicate paleoenvironments ranging from below storm wave base to upper shoreface. The edgewise, mounded, and parallel-bedded dolomite-chip breccia indicates slope failure and reworking of the lower shoreface during large storms. Facies relationships allow generalized division of these strata into three units. The lowermost unit, Unit A, comprises intimately interbedded laminated dolomite (“cap dolostone”), dolomite-chip breccia, and sandstone facies and is 17 m thick. Unit A apparently grades upward into Unit B, a 45-m-thick interval of the sandstone facies. Unit C, 28 m thick, rests sharply on Unit B, and comprises a basal laminated dolomite facies and the limestone and argillite facies. Units A and B may indicate a regressive wave-dominated coast that was influenced by large storms (highstand systems tract). Unit C indicates near storm wave base deposition and an overall deepening, as shown by dark argillite beds of the overlying upper member of the Pocatello Formation (transgressive systems tract). δ13C and δ18O values from dolomite and limestone samples of Units A and C are similar to values from local, regional, and transglobal cap carbonate intervals. δ13C values range from −1.9 to −5.6‰ and δ18O values range from −10.2 to −17.4‰, with no systematic correlation between C- and O-isotope values. δ13C values are consistent with previously reported values from the Pocatello Formation and are similar to

Dehler, C.M., Anderson, K., and Nagy, R., 2011, New descriptions of the cap dolostone and associated strata, Neoproterozoic Pocatello Formation, southeastern Idaho, USA, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 183–194, doi:10.1130/2011.0021(08). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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Dehler et al. values from the alleged Marinoan Noonday Formation in Death Valley, California, and the Marinoan Maieberg Formation in Namibia. Collective data from the cap dolostone and associated strata of the Pocatello Formation suggest protracted mixed siliciclastic-carbonate deposition on a stormdominated shelf at ca. 665 Ma.

INTRODUCTION Neoproterozoic cap carbonate units, which overlie glacial and subaqueous mass flow deposits worldwide, are interpreted to indicate the recovery from glacial conditions of controversial severity (e.g., Hoffman et al., 1998; Fairchild and Kennedy, 2007; Allen and Etienne, 2008; LeHeron et al., 2011). Cap dolostones (the basal transgressive tract of the cap carbonate unit) with distinctive lithologic and geochemical similarities to the type Marinoan carbonate (Keilberg member, Maiberg Formation, Namibia; Hoffman, 2011) are presumed to be 635 Ma everywhere and thus are used to mark the base of the Ediacaran Period of the latest Neoproterozoic (Condon et al., 2005; Knoll et al., 2006). The Marinoan cap dolostone is typically pink to tan, has an average global thickness of 18.5 m, and may contain some or all of the following, depending on water depth: paralleland (or) cross-laminated microcrystalline dolomite, peloids, geoplumb stromatolites and(or) tubestone, sheet cracks, giant wave ripples, teepee structures, barite and aragonite crystal fans, and negative δ13C values that decrease upward (e.g., Hoffman et al., 2007; Hoffman, 2011). “Sturtian” cap carbonate units, which are interpreted to predate the Marinoan glaciation and range in age from 660 to 720 Ma (Hoffman and Li, 2009), are less well studied. These cap units are organic-rich, dark gray, iron-rich, contain rhythmic and anastomosing laminae, and have δ13C values >0 (Kennedy et al., 1998). Despite the importance of cap carbonate units in understanding deep-time paleoclimate and their potential utility as time lines, few alleged “Sturtian” cap units have been well described and almost no cap units have been placed in a solid geochronologic framework. Previous work on the cap dolostone and associated strata of the Pocatello Formation indicates that they represent deposition at ca. 665 Ma (Fanning and Link, 2004, 2008; Dehler et al., 2009), yet the lithologic and isotopic character of the strata are similar to alleged Marinoan cap carbonate units (Lorentz et al., 2004; Kirkham et al., 2009; Anderson and Dehler, 2010), only one of which has been directly dated at 635 Ma (Doushantuo Formation; Condon et al., 2005). This discrepancy has implications for our understanding of the number and nature of post-glacial alkalinity events, and for testing the snowball Earth hypothesis (Hoffman et al., 1998). The cap dolostone and associated strata in the Pocatello Formation are important because they represent one of few cap carbonate units in western Laurentia and one of the only cap intervals with geochronologic control. Until now, most information on the Pocatello cap interval has come from two localities, north and immediately south of the Portneuf Narrows (also known as

Portneuf Gap), Idaho (Locations 1 and 2, Fig. 1). However, the sections at these two localities are incomplete. The northern section lacks a cap dolostone, and the southern section, although it does exhibit a 1-m-thick cap dolostone, is only ~10 m thick and is poorly exposed. Herein we present preliminary facies, stratigraphic, and stable isotopic data from newly discovered exposures of the cap dolostone and associated strata of the Pocatello Formation, south of Portneuf Gap in the Fort Hall area (Locations 3 and 4, Fig. 1). The results from these new measured sections modify previous stratigraphic and sedimentologic interpretations, which are key to understanding this cap carbonate unit and how it relates to others. Scout Mountain Member of the Pocatello Formation The Pocatello Formation is a >1.5-km-thick mixed siliciclastic, carbonate, and volcanic unit that is exposed for 100 km along strike between the Pocatello area of southern Idaho, south to the Utah portion of Cache Valley (Fig. 1). The formation comprises 3 members, has no exposed base, and is overlain by the Blackrock Canyon Limestone (Link, 1983). The lowest member, the Bannock Volcanic Member (200–450 m thick), is composed of mafic metavolcanic rocks (pillow lava, agglomerate, and greenstone) and volcaniclastic rocks (Link et al., 1994; Keeley and Link, this volume). This unit grades upward into the Scout Mountain Member (~<800 m thick), which contains a wide variety of lithotypes including quartz, arkosic, and lithic arenite, argillite, clast-supported orthoconglomerate, intra- and extra-basinal diamictite, and subordinate limestone, and dolostone. This unit grades upward into the informal upper member of the formation, which comprises (>600 m thick) of laminated argillite with thin sandy beds. The Pocatello Formation is interpreted to represent one of many rift basins that were fed by glaciers along the western Cordilleran margin in late Neoproterozoic time (Stewart, 1972; Link, 1982). The Bannock Volcanic Member mafic volcanic rocks are thoeliitic-alkaline to alkaline in composition and indicate intraplate volcanism, which is consistent with a rift basin setting (Harper and Link, 1986). The Scout Mountain Member records rapid deposition of immature sediments, including interpreted glacial and resedimented glacial deposits. Evidence for a glacial origin is limited, but the lateral extent of the deposits (100 km), rare striated clasts, and correlation with glacial deposits along the Cordilleran margin argue for a glacial origin (e.g., Crittenden et al., 1983; Link et al., 1994). The carbonate and sandstone units in the upper Scout Mountain Member indicate tidal, shoreface, and

Cap dolostone and associated strata, Neoproterozoic Pocatello Formation

1

2 SYMBOLS Contact: dashed where approximate,dotted where concealed Normal fault: dashed where approximate,dotted where concealed Anticline axial trace: dashed where approximate,dotted where concealed Overturned anticline axial trace : dashed where approximate,dotted where concealed

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Strike and dip of bedding Strike and dip of overturned bedding

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Strike and dip of vertical bedding measured section Section 1 Section 2 Section 3 Section 4

666 +/- 6 Ma

North Gap South Gap North Fort Hall Fort Hall

0.5 mile

MAP UNITS Q T

Quaternary units Tertiary units Neoproterozoic units

usl ud

Zc Zb

Zpub Zpua

Caddy Canyon Quartzite Blackrock Canyon Limestone Pocatello Formation upper member unit B--quartzite unit A--phyllite

ls3 ls2 ls1 ld Zpb

Pocatello Formation Scout Mountain Member upper sandstone and limestone unit upper diamictite lower sandstone unit 3 lower sandstone unit 2 lower sandstone unit 1 lower diamictite Pocatello Formation Bannock Volcanic Member

Figure 1. Geologic map showing distribution of Pocatello Formation members in the Portneuf Gap area. Measured sections 1 through 4 are shown with red lines. The “upper sandstone and limestone unit” of the Scout Mountain Member is the focus of this study. Map modified from Rodgers et al. (2006).

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deeper water conditions (Link, 1983). The upper member of the Pocatello Formation is interpreted to indicate an overall deepening caused by post-glacial eustatic rise (Link et al., 1994). CAP DOLOSTONE AND ASSOCIATED STRATA OF THE FORT HALL AREA

upper member

What has traditionally been called the cap dolostone of the Pocatello Formation is a 1-m-thick “pink dolostone,” which sits

Brigham Group

section of study

<~665 Ma

Scout Mountain Member

“Sturtian”

Cryogenian

Blackrock Cyn. Limestone

? M

580 Ma

Pocatello Fm.

Ediacaran

C

1 km

in sharp contact on a prominent extrabasinal diamictite known as the “upper diamictite” and is located south of the Portneuf Gap (Link, 1982, terminology; Section 2 in Figs. 2, 3, and 4). Newly explored outcrops along strike of this cap dolostone to the south reveal a thicker section of cap dolostone, which is intimately interbedded with dolomite-chip breccia and sandstone (Sections 3 and 4 in Figs. 2, 3, and 4). North of Portneuf Gap, the cap dolostone is absent and instead, a meter-thick “dolomite-chip breccia” lies directly on the upper diamictite (Link, 1982, terminology).

Fe

<~680 Ma

base not exposed

limestone dolomite

100 m

KEY shale, siltstone

Bannock Volc. Mbr.

Neoproterozoic section, Pocatello, Idaho

<~705 Ma

diamictite conglomerate sandstone mafic volcanics

Fe

iron-rich interval sequence boundary

M

Marinoan

Pocatello Formation Portneuf Gap

Figure 2. Stratigraphic column showing upper Neoproterozoic (Cryogenian and Ediacaran) strata of southeastern Idaho (left) and generalized stratigraphy of the Pocatello Formation, and location of the cap-carbonate sequence in the upper Scout Mountain Member. Modified from Fanning and Link (2004). The 705 Ma maximum depositional age is reported from Keeley and Link (this volume). The 680 and 665 Ma maximum depositional ages are reported from Dehler et al. (2009).

Scout Mountain Member

upper member

North Gap Section 1

original ‘pink dolostone’

xxx 667 +/- 5 Ma

South Gap Section 2

xxx reworked tuff dz detrital zircon sample SB sequence boundary

diamictite

reworked zone

acritarchs

North Fort Hall Section 3

crinkly bedding flute casts

stromatolites

plane beds intraclasts crossbedding aragonite fans (replaced)

slump fold

limestone and argillite interbedded argillite w/ sandstone interbeds

dolomite breccia

hummocky cross stratification convolute bedding

peloidal dolomite

sandstone

wave ripples low -angle and horizontal laminations groove casts

?

Fort Hall Section 4

A

B

C

Units

dz ca. 680 Ma

dz 666 +/- 6 Ma

Scout Mountain Member

upper member

Figure 3. Stratigraphic correlation of measured sections across the Portneuf Gap area. See Figure 1 for location of measured sections. Note the changes in facies across the N-S transect. The Fort Hall and North Fort Hall localities have only recently been discovered and help to clarify facies relationships observed in the two original study localities, North and South Gap.

SB

10 m

deepening upwards shallowing upwards

Explanation

Cap dolostone and associated strata, Neoproterozoic Pocatello Formation 187

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Units

C 60

FACIES sandstone (upper shoreface (and foreshore?) dolomite-chip breccia (lower shoreface-offshore slope failure)

55 40

laminated dolomite units (upper shoreface to offshore) silty sandstone (reworked diamictite matrix)

B 35 25

diamictite (glacial marine and mass flow) break in section 20 19 18

18

17

17

16

16 6

16 15

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15

14

14

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1 3 13

14 13 12 11 10 9

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7

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3

3

2

2

2

1

original ‘pink 2 dolostone’ 1

1

1

0

0

0

0

South Gap Section 2 North

North Fort Hall Section 3

A

Fort Hall Section 4 South

Figure 4. Examples of lateral and vertical facies changes and depositional environments of the basal cap carbonate interval in the South Gap area. Note the lateral discontinuity of different carbonate facies. See Figure 1 for location of measured sections.

Cap dolostone and associated strata, Neoproterozoic Pocatello Formation Cap-Carbonate Interval Geochronology Zircon U-Pb geochronology from volcaniclastic units in the cap dolostone and associated strata indicates deposition ≤665 Ma. A green reworked tuffaceous siltstone with abundant euhedral, zoned igneous zircons lies near the top of Unit B (~45 m from base of interval, North Gap locality; Figs. 1, 2, and 3) and yields a U-Pb age of 667 ± 5 (sensitive high-resolution ion microprobe [SHRIMP] concordia age; Fanning and Link, 2004, 2008). Igneous zircon grains, also euhedral and zoned, from finegrained sandstone interbeds in the basal cap dolostone interval, Unit A (4 m above base of interval of the Fort Hall section; Figs. 2 and 3) yield a similar age of 666 ± 6 Ma (SHRIMP concordia age; Dehler et al., 2009), which provides the understanding that the cap dolostone and associated strata are ca. 665 Ma or younger. The upper diamictite matrix immediately below the cap dolostone interval at the Fort Hall locality yields a maximum depositional age of ca. 680 Ma, as does an arkosic sandstone unit ~110 m from the base of the Scout Mountain Member (SHRIMP concordia ages on youngest populations; Fig. 2; Dehler et al., 2009). The basal Scout Mountain Member yields a maximum depositional age of 705 ± 5 Ma (SHRIMP concordia age; Keeley and Link, this volume). Regionally, the upper age constraint on the cap dolostone and associated strata is a ca. 580 Ma Rb-Sr date on the Browns Hole Formation in northern Utah, which correlates with local strata that is ~2 km upsection (Christie-Blick and Levy, 1989). The young zircon populations in the cap dolostone and associated strata are interpreted to have been derived from approximately syndepositional volcanism and thus likely very closely bracket maximum depositional ages. This interpretation is based on the progressively younger zircon ages found upsection in the Pocatello Formation, lack of young zircons in overlying formations, and correlation to other dated post-glacial deposits. The maximum depositional age of the lowermost Scout Mountain Member is 705 Ma ± 5 Ma (Keeley and Link, this volume). The maximum depositional age for the diamictite and lower units in the overlying Scout Mountain Member is ca. 680 Ma and the maximum depositional age of the cap dolostone and associated strata, Units A and B, in the upper Scout Mountain Member is 666 Ma ± 6. This younging of ages (from 705 Ma to 680 Ma to 666 Ma) would be more likely preserved if volcanism occurred as sediments were being deposited. Young zircon grains are stratigraphically constrained to the Pocatello Formation. Detrital zircon geochronologic analyses conducted on overlying Neoproterozoic and Cambrian units in the region show a dearth of 705–665 Ma grain populations (Stewart et al., 2001; Keeley et al., 2009; Reed et al., 2010; Link et al., 2011). The 666 Ma date of the cap dolostone interval from the Pocatello Formation also overlaps in age with two other post-glacial deposits. The postglacial black shale of the Aralka Formation, Australia is 657 ± 5 Ma (Re-Os on black shale, Kendall et al., 2009), and the cap carbonate in the Datangpo Formation, south China is 663 ± 4 Ma (U-Pb zircons in ash, Zhou et al., 2004). On the other hand, as

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described below, the cap dolostone and associated strata share sedimentological and isotopic similarities with Marinoan and alleged Marinoan cap-carbonate intervals worldwide, most of which, however, lack good geochronologic control. Facies Analysis and Paleoenvironmental Interpretation Seven facies are described herein and include diamictite, silty sandstone (reworked diamictite matrix), laminated dolomite, dolomite and sandstone, sandstone, dolomite-chip breccia, and argillite and limestone facies (Table 1). The upper diamictite in the newly described sections (Localities 3 and 4, Figs. 1 and 2) was only examined immediately below the contact with the overlying cap interval. The chaotic fabric of the clasts and lack of glacial indicators suggests subaqueous mass flow deposition, although the sediments were likely derived from glaciers (Link, 1982; Crittenden et al., 1983). The water depth is difficult to discern due to a lack of sedimentary structures, but this facies rests on laminated argillite that represents deep-water deposition. The silty sandstone facies rests sharply on the diamictite and grades upward into the laminated dolomite facies. The silty sandstone facies takes on the same color as the underlying matrix of the diamictite, is a calcareous silty sandstone with granules and pebbles of dolomite, and shows local symmetrical ripplemarks and laminations. This facies is similar sedimentologically to the overlying laminated dolomite facies (Table 1) and is interpreted to represent reworking of the diamictite during initiation of cap dolostone deposition. Although it is not yet known how much time might be missing on this reworked surface, it represents reworking of the underlying diamictite matrix and mixing with the dolomite facies. The laminated dolomite (“pink dolostone” of Link, 1982) is very thinly bedded and becomes thinly bedded and grades into the dolomite and sandstone facies upsection on a meter scale (Fig. 4). Sedimentary structures in the laminated dolomite facies indicate deposition below storm wave base where it is purely dolomitic, to deposition near and above fair weather wave base where it is interbedded with rippled sandstone with rare hummocky cross stratification (interbedded dolomite and sandstone facies). Link (1982) interpreted this facies to indicate deposition in a tidal flat, but did not discount the potential of a deep-water origin. Within Unit A, this facies is considered to be the cap dolostone (Figs. 3 and 4). A similar dolomite facies appears at the base of Unit 3, yet is not considered part of the cap dolostone because it is separated by a potential bounding surface, lacks abundant lamination, and has different weathering characteristics (Figs. 3 and 4). The dolomite-chip breccia facies occurs twice in meter-thick, laterally discontinuous beds, and can be traced laterally into the sandstone and dolomite facies (Fig. 4). Breccia clasts are composed of laminated dolomite and the matrix is sandstone, both of which are directly sourced from the sandstone and dolomite facies. Breccia clasts are platy and range in size from a few centimeters to a meter (long axis). Some of the breccia shows chaotically oriented

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Facies Diamictite

Dehler et al. TABLE 1. FACIES WITHIN AND BELOW THE CAP-CARBONATE INTERVAL IN THE PORTNEUF GAP AREA Description Interpretation Matrix: purple-gray, silt to coarse sand, poorly sorted. Glaciomarine and reworked glacial Clasts: granule to boulder, dominantly quartzite with lesser volcanic clasts. deposits (Link, 1982).

Silty sandstone facies

Dm-thick interval of calcareous siltstone, sand and granules. Purple-gray color, thin bedding, normal grading, sinuous symmetric and interference ripplemarks.

Reworked diamictite matrix.

Laminated dolomite

Tan-pink color, thin bedding with internal lamination, peloidal dolomicrite to dolosiltite. Mud-draped symmetric and interference ripples, rare hummocky cross-stratification.

Lower shoreface to offshore, near or below fair-weather wave base, rare combined flow, density currents, microbial influence, suspension settling.

Interbedded dolomite and sandstone

Laminated dolomite facies rhythmically interbedded with fine-grained sandstone in thin to medium beds. Amalgamated and solitary hummocky cross-stratification (<20 cm high, 1 m wide), rare dolomite-chip breccia in hummocks, rare m-thick channel forms, symmetric and interference ripples, positive flute casts.

Upper to lower shoreface, near fairweather wave base, combined flow, density currents.

Sandstone

Fine- to medium-grained sand, planar medium to very thick beds. Rare trough crossbedding. Small dolomite intraclasts (avg. diameter 3 cm) at base of some beds. Possible swaley beds. Few beds of green argillite and volcaniclastic material. Load, flute, groove casts, dish structures, parting lineations. Channelized base at North Fort Hall locality. Organic-walled microfossils.

Upper shoreface to foreshore or beach?

Dolomite-chip breccia

Thickly bedded. Clast-supported. Clasts composed of laminated dolomite facies, typically tabular (<1 m on long axis), syndepositional deformation. Fabric types: chaotic, fan-shaped, bedding-parallel. Scours, rare onlapping relationships, breccia mounds.

Deposits from large storms, combined flow, mass flow, gravity slide processes causing failure of lower shoreface.

Argillite and limestone facies

Meter-thick intervals of green and brown argillite interbedded with medium beds of bluish-gray limestone. Horizontal laminations, climbing ripples. Minor quartz silt component. “Ghost” aragonite fans (<3 cm tall).

Relative deepest water. Offshore. Suspension settling and crystal fan growth.

clasts including edgewise clasts, rosettes of clasts, and breccia mounds. Another subfacies of breccia has bedding parallel clasts in a sandy matrix, indicating reworking. Some of the clasts are bent to folded, indicating they were soft during transport. The dolomite-chip breccia facies is very similar to intraformational breccia units of the Upper Cambrian to Lower Ordovician Snowy Range Formation, Wyoming (Myrow et al., 2004) and the Lower Cambrian Sellick Hill Formation, South Australia (Mount and Kidder, 1993), which were both interpreted to indicate mass flow and gravity slide processes on the lower shoreface, caused by large storms. Mount and Kidder (1993) suggest that the storms must have been very extreme to produce the observed chaotic fabrics. At the North Gap locality, the dolomitechip breccia facies is exposed in a solitary 1-m-thick interval and rests directly on the diamictite. Link (1982) interpreted the breccia here to indicate a transgressive lag and that it was unconformable with the laminated dolostone (in the cap dolostone interval). Now that these two facies are known to intercalate (Fig. 4), no unconformity is necessary between the two facies. The sandstone facies is a dominantly massive subfeldspathic sandstone that is crudely and thickly bedded. Rare groove casts, plane beds, parting lineations, and trough crossbedding are present and, considering association with other facies, likely represent the shoreface/foreshore environment. Link (1987) also interpreted this facies to indicate a shallow marine to beach setting. At

the North Fort Hall locality, the basal sandstone has a channelized base. Trough crossbedding and rip ups of laminated dolomite are also observed at the base (Fig. 4). This facies has a few decimeter-scale interbeds of green argillite, some of which have yielded organic-walled microfossils (Fig. 5). The argillite and limestone facies consists of covered intervals with subcrops of limestone and argillite. The argillite is green to brown and laminated. The limestone is bluish gray and shows laminations, replaced aragonite fans, and climbing ripples. The replaced aragonite fans are continuous on a decimeter scale and range in height from 2 mm to 2 cm. Lorentz et al. (2004) discovered replaced aragonite fans in this same facies and stratigraphic interval north of the Portneuf Narrows (Section 1, North Gap Section, Figs. 1 and 3). Because there is no indication of exposure and this facies is overlain by a significant thickness of argillite of similar character, it likely indicates deeper water deposition. Other workers have interpreted aragonite fans in Neoproterozoic deposits to indicate deep water (and maximum flooding) (e.g., Hoffman, 2011). Stratigraphy An ~90-m-thick interval of strata rests sharply on the upper diamictite in the Fort Hall area south of Portneuf Gap. This interval is mapped as the “upper sandstone and limestone unit” of the

Cap dolostone and associated strata, Neoproterozoic Pocatello Formation upper Scout Mountain Member on the Inkom 7.5′ quadrangle (Fig. 1; Rodgers et al., 2006). This interval is the focus of this study because it not only captures lateral and vertical facies changes of the cap dolostone, it also includes dolomite and limestone units higher in the section that may be related to the underlying cap dolostone unit. The top of the interval of study coincides with the contact between the Scout Mountain Member and the overlying upper member of the Pocatello Formation (Figs. 1, 2, and 3). The upper member is a relatively homogeneous argillaceous and sandy unit that is hundreds of meters thick (Link, 1982). Facies relationships allow generalized division of these strata into three units. The lowermost unit, Unit A, comprises intimately interbedded laminated dolomite (cap dolostone), dolomite-chip breccia, and sandstone facies and is ~17 m thick. The relationship between the laminated dolomite and the other facies indicates dynamic mixed siliciclastic-carbonate deposition that was alternating between below fair weather wave base to below storm wave base. The facies indicating deeper water

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(laminated dolomite and sandstone and dolomite facies) are more common in the lower part of Unit A, whereas sandstone with hummocky cross-stratified sandstone is higher in Unit A, indicating an overall shallowing upward trend (Figs. 3 and 4). Unit B is apparently gradational with underlying unit A and is a 45-m-thick interval of the sandstone facies. The sandstone within this unit is similar to the sandstone beds below except it is significantly thicker, relatively homogeneous, and it is massive for the most part. It is also different in that it contains sedimentary structures such as soft sediment deformation features, parting lineations, plane beds, faint rare crossbeds, and rare dolomite-chip lags. These features indicate the likely shoreface part of the wave-dominated system, whereas the interbedded sandstone facies below likely indicates deposition below fair weather wave base. Unit C, 28 m thick, rests sharply on Unit B, and begins with a basal laminated dolostone facies that is 6 m thick. Above this, interbedded argillite, some sandstone, and medium beds of

Figure 5. Representative organic-walled microfossils (“acritarchs”) from green argillite of the Scout Mountain Member of the Pocatello Formation (sample SMM-08-02), including: spine-bearing acritarchs (A, B), textured vesicles (C), and a smaller specimen comprising many minute vesicles (cf. Bavlinella faveolata) (D). Fossils were recovered via standard hydrofluoric acid maceration techniques. Although these specimens have not yet been assigned a taxonomic identity, they likely represent marine eukaryotes (e.g., Vidal and Knoll, 1983; Servais, 1996). Scale bar is 10 microns for A–D.

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limestone are present. This unit was only measured up to the last exposed limestone bed, which also marks the top of the Scout Mountain Member (Link, 1982; Rodgers et al., 2006). The change back to carbonate and argillite indicates an overall deepening and likely marks a flooding surface. Correlation of the Fort Hall section with the other sections (Figs. 3 and 4) shows Unit A becoming more sandy toward the north and probably thinning below Unit B to a few meters at the North Gap locality. Unit B appears to maintain about the same thickness between the southernmost and the northernmost section. The basal dolomite unit of Unit C is not present in the North Gap section and limestone is more abundant (Fig. 4).

The range of values reported herein is similar to those reported from the Noonday Formation of Death Valley, southeastern California, and from the Maiberg Formation of Namibia. The Sentinel Peak (cap dolostone) and lower Radcliff members of the Noonday Dolomite show a range of δ13C values from −1.8 to about –6.1‰ and δ18O values of −5 to about −15‰ (Petterson et al., 2011). Cap carbonate δ13C values from the Maieberg Formation in Namibia range from −3 to −5‰ and δ18O values are between −5‰ and −10‰ (Hoffman et al., 1998). The range of δ18O values from both the Pocatello Formation and the Noonday Dolomite are relatively low and could be an indication of meteoric diagenesis or metamorphism (e.g., Kaufman and Knoll, 1995).

Stable Isotope Analyses

DISCUSSION

Dolomite and limestone samples from Units A and C were microdrilled to obtain powder for carbon- and oxygen isotope analyses. Methods of sample preparation and analyses are those followed in Dehler et al. (2005). Resulting δ13C values range from −1.9 to −5.6‰ and δ18O values range from −10.2 to −17.4 ‰. A crossplot of these values from the South Gap localities shows overlap with values previously reported by Lorentz et al. (2004) from the North Gap area, yet our values are typically more positive (Fig. 6). This difference in values is likely due to the fact that the majority of their samples came from Unit 3 equivalent in the North Gap area (Fig. 3), which has been partially altered to marble, and therefore, those values would be expected to partly reflect metamorphism and show more negative values (e.g., Kaufman and Knoll, 1995). Values are also consistent with previous values reported by Smith et al. (1994).

If there were no age constraints on the Pocatello cap dolostone and associated strata and no other reported ca. 660 Ma dates from other diamictite-cap carbonate bearing successions (Zhou et al. 2004; Fanning and Link, 2008; Kendall et al., 2009), the newly described sections would likely be considered 635 Ma in age because of their similarities to the Namibian and Doushantuo cap carbonate successions and other alleged Marinoan successions. However, the available geochronology, in combination with the geologic context, suggests that the Pocatello strata were deposited ca. 665 Ma. This interpretation has implications for using cap carbonate units as timelines (Halverson, 2006; Corsetti and Lorentz, 2006) and brings into question the use of the Marinoan cap carbonate as the global stratotype section and point of the basal Ediacaran (Knoll et al., 2006). Unless the geochronologic context of a cap-carbonate unit is well constrained, or until

Figure 6. Isotope crossplot comparing carbon and oxygen isotope values from the South Gap area (this study) and the North Gap area (Lorentz et al., 2004). Note that the North Gap locality values are relatively more negative than those at the South Gap locality and are probably reflecting diagenetic alteration or metamorphism of the North Gap carbonates.

Cap dolostone and associated strata, Neoproterozoic Pocatello Formation cap-carbonate units can be demonstrated to be coeval, caution must be taken in assuming synchroneity of deposition. The sedimentology of the basal Pocatello cap dolostone unit indicates large storm events. In the basal (Marinoan) cap sequences on other continents, sedimentary structures interpreted to be giant wave ripples may also indicate infrequent storms with sustained high velocity winds (Allen and Hoffman, 2005). Interpreted storm features in the Pocatello Formation may be a different expression of the same type of post-glacial process. The Pocatello Formation cap dolostone unit, if ca. 665 Ma, indicates that there were at least two paleoclimate episodes (665 Ma, 635 Ma) that experienced similar post-glacial processes (e.g., large storms) resulting in geologically similar sequences. Most diamictite-cap-carbonate contacts described in the literature are interpreted to be gradational, despite many of them having sharp contacts (e.g., Fairchild and Kennedy, 2007). The Pocatello Formation diamictite-cap dolostone contact appears to be unconformable; it is sharp with a reworked zone of diamictite matrix above the contact. The presence of an unconformity is supported by a jump in young detrital U-Pb age populations across the contact, from ca. 680 Ma below, to ca. 666 above (Fanning and Link, 2004, 2008; Dehler et al., 2009; Keeley and Link, this volume). This relationship may have implications for our understanding of post-glacial processes. Facies associations between the laminated dolomite of Unit A (cap dolostone) and other facies including the dolomite-chip breccia, sandstone and dolomite, and sandstone facies indicate that cap-carbonate deposition was not laterally continuous on a basinal scale, but rather these carbonates were deposited in a dynamic environment, mixing with siliciclastic sediment. Furthermore, Units A through C show negative C-isotope values like cap-carbonate units elsewhere, yet likely record at least one bounding surface. These points suggest that the cap dolostone and related strata of the Pocatello Formation record a protracted and dynamic history of mixed siliciclastic-carbonate deposition (cf. Kennedy and Christie-Blick, 2011). CONCLUSIONS Lithofacies, paleoenvironmental interpretation, and isotopic values are similar to other cap intervals that are dated as, or interpreted to be, Marinoan (635 Ma) in age, yet there are many arguments for the Pocatello cap interval to be ca. 665 Ma. Observations from this interval contrast with generalizations about cap carbonate successions that include: (1) the contact between the underlying diamictite unit and the cap dolostone may be an unconformity; (2) the cap dolostone interfingers with several other facies, including siliciclastic facies; and (3) there may be at least one bounding surface within the overall post-diamictite interval. Because other post-glacial strata date to ca. 660 Ma, it may be that this is a 665 Ma cap carbonate interval that looks very similar to Marinoan cap carbonate intervals. These findings have implications for using cap carbonates as global marker units and for interpreting them to indicate global alkalinity events.

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DIRECTIONS TO CAP SEQUENCE EXPOSURES From Interstate Highway 15, northbound or southbound, take the Fort Hall Mine/Landfill exit (Exit 63), ~3 mi southeast of Pocatello. Head south on Fort Hall Mine Road. Cross the railroad tracks and the Portneuf River. Continue for 0.2 mi to Portneuf Road. Cross this road and drive straight for 0.6 mi. At this point, take a left onto the dirt road (before the Fort Hall Mine Landfill entrance). Proceed along this road. It will curve to the south, up Fort Hall Mine Canyon, and will turn into an all-terrain vehicle track (1 mi). Park here and walk up road that leads south up to Fort Hall Mine to the east, until you are above the mine. Here will be a faulted section of the cobble conglomerate in the Scout Mountain Member. Walk up the ridge and, once you reach the diamictite, veer to the north toward the drainage. To see the units below in this member, walk up next drainage to north to find the “lower diamictite” (correlative with Oxford Ridge Diamictite (see Keeley and Link, this volume). The upper diamictite forms a cliff near the top of the ridge. Sharply overlying the diamictite is the cap carbonate interval. The climb is about ~800 vertical feet total. Walk exposures along strike to the north to next ridge to see lateral facies changes in the cap sequence. ACKNOWLEDGMENTS This research was supported by National Science Foundation grant EAR-0819759. Field work was provided by Scott Sallay, Hans Anderson, Kelly Bradbury, Dawn Hayes, Esther Kingsbury, Liz Petrie, Dan Rybczynski, Eva Lyon, and Sienna and other field dogs. Ryan Petterson and Maya Elrick provided helpful reviews. Mark Fanning, Paul Link, and Adolph Yonkee provided helpful comments. REFERENCES CITED Allen, P.A., and Etienne, J.L, 2008, Sedimentary Challenge to Snowball Earth: Nature Geoscience, v. 1, p. 817–825. Allen, P.A., and Hoffman, P.F., 2005, Extreme winds and waves in the aftermath of a Neoproterozoic glaciation: Nature, v. 433, p. 123–127, doi:10.1038/ nature03176. Anderson, K.R., and Dehler, C.M., 2010, Looks are deceiving: Stable isotope data from the Pocatello Formation cap-carbonate sequence and its implications for cap-carbonate correlations: Geological Society of America Abstracts with Programs, v. 42, no. 5, p. 299. Christie-Blick, N., and Levy, M., 1989, Concepts of sequence stratigraphy, with examples from strata of Late Proterozoic and Cambrian age in the western United States in Christie-Blick, N., and Levy, M., eds., Late Proterozoic and Cambrian Tectonics, Sedimentation, and Record of Metazoan Radiation in the Western United States: Washington, D.C., American Geophysical Union, 28th International Geological Congress Field Trip Guidebook T331, p. 23–38. Condon, D., Zhu, M., Bowring, S., Wang, W., Yang, A., and Jin, Y., 2005, U-Pb Ages from the Neoproterozoic Doushantuo Formation, China: Science, v. 308, p. 95–98, doi:10.1126/science.1107765. Corsetti, F.A., and Lorentz, N.J., 2006, On Neoproterozoic cap carbonates as chronostratigraphic markers, in Xiao, S., and Kaufman, A.J., eds., Neoproterozoic Geobiology and Paleobiology, p. 273–294. Crittenden, M.D., Jr., Christie-Blick, N., and Link, P.K., 1983, Evidence for two pulses of glaciation during the Late Proterozoic in northern Utah and southeastern Idaho: Geological Society of America Bulletin, v. 94, p. 437–450, doi:10.1130/0016-7606(1983)94<437:EFTPOG>2.0.CO;2.

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Dehler, C.M., Elrick, M.E., Bloch, J.D., Karlstrom, K.E., Crossey, L.J., and Des Marais, D., 2005, High-resolution δ13C stratigraphy of the Chuar Group (ca. 770–742 Ma), Grand Canyon: Implications for mid-Neoproterozoic climate change: Geological Society of America Bulletin, v. 117, no. 1/2, p. 32–45, doi:10.1130/B25471.1. Dehler, C.M., Fanning, C.M., and Link, P.K., 2009, Getting better with age: New U-Pb SHRIMP data from the “Sturtian” Scout Mountain diamictitecap sequence, Pocatello Formation, Idaho: Geological Society of America Abstracts with Programs, v. 41, no. 7, p. 591. Fairchild, I.J., and Kennedy, M.J., 2007, Neoproterozoic glaciation in the Earth System: Journal of the Geological Society, v. 164, p. 895–921, doi:10.1144/0016-76492006-191. Fanning, C.M., and Link, P.K., 2004, U-Pb SHRIMP ages of Neoproterozoic (Sturtian) glaciogenic Pocatello Formation, southeastern Idaho: Geology, v. 32, p. 881–884, doi:10.1130/G20609.1. Fanning, C.M., and Link, P.K., 2008, Age constraints for the Sturtian Glaciation; data for the Adelaide Geosyncline, South Australia and Pocatello Formation, Idaho, USA: Selwyn Symposium 2008 GSA Victoria Division, Abstracts 91, p. 57–61. Halverson, G.P., 2006, A Neoproterozoic Chronology in Xiao, S., and Kaufman, A.J., eds., Neoproterozoic Geobiology and Paleobiology, p. 231–272. Harper, G.D., and Link, P.K., 1986, Geochemistry of Upper Proterozoic riftrelated volcanics, northern Utah and southeastern Idaho: Geology, v. 14, p. 864–867, doi:10.1130/0091-7613(1986)14<864:GOUPRV>2.0.CO;2. Hoffman, P.F., 2011, Strange bedfellows: glacial diamictite and cap carbonate from the Marinoan (635 Ma) glaciation in Namibia: Sedimentology, v. 58, p. 57–119, doi:10.1111/j.1365-3091.2010.01206.x. Hoffman, P.F., and Li, Z.X., 2009, A paleogeographic context for Neoproterozoic glaciation. Paleogeography, Paleoclimatology, Paleoecology, doi: 10.1016/j.palaeo.2009.03.013. Hoffman, P.F., Kaufman, A.J., Halverson, G.P., and Schrag, D.P., 1998, A Neoproterozoic Snowball Earth: Science, v. 281, p. 1342–1346, doi:10.1126/ science.281.5381.1342. Hoffman, P.F., Halverson, G.P., Domack, E.W., Husson, J.M., Higgins, J.A., and Schrag, D.P., 2007, Are basal Ediacaran (635 Ma) post-glacial “cap dolostones” diachronous?: Earth and Planetary Science Letters, v. 258, p. 114–131, doi:10.1016/j.epsl.2007.03.032. Kaufman, A.J., and Knoll, A.H., 1995, Neoproterozoic variations in the C-isotopic composition of seawater: stratigraphic and biogeochemical implications: Precambrian Research, v. 73, p. 27–49, doi:10.1016/0301 -9268(94)00070-8. Keeley, J.A., and Link, P.K., 2011, this volume, Middle Cryogenian (“Sturtian”) Pocatello Formation: Field relations on Oxford Mountain and the Portneuf area, southeast Idaho, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, doi:10.1130/2011.0021(07). Keeley, J., Link, P.K., and Dehler, C.M., 2009, Detrital Zircon Analysis of the Neoproterozoic to Cambrian Brigham Group in the Southern Portneuf Range, Southeast Idaho: Geological Society of America Abstracts with Programs, v. 41, no. 6, p. 45. Kendall, B., Creaser, R.A., Calver, C.R., Raub, T.D., and Evans, D.A.D., 2009, Correlation of Sturtian diamictite successions in southern Australia and northwestern Tasmania by Re-Os black shale geochronology and the ambiguity of “Sturtian”-type diamictite-cap carbonate pairs as chronostratigraphic marker horizons: Precambrian Research, v. 172, p. 301– 310, doi:10.1016/j.precamres.2009.05.001. Kennedy, M.J., and Christie-Blick, N., 2011, Condensation origin for Neoproterozoic cap carbonates during deglaciation: Geology, v. 39, p. 319–322, doi:10.1130/G31348.1. Kennedy, M.J., Runnegar, B., Prave, A.R., Hoffman, K.-H., and Arthur, M.A., 1998, Two or four Neoproterozoic glaciations?: Geology, v. 26, p. 1059– 1063, doi:10.1130/0091-7613(1998)026<1059:TOFNG>2.3.CO;2. Kirkham, K., Dehler, C.M., and Link, P.K., 2009, Investigating a Sturtian cap-carbonate sequence, Scout Mountain Member, Pocatello Formation, Idaho: Geological Society of America Abstracts with Programs, v. 41, no. 6, p. 45. Knoll, A.H., Walter, M.R., Narbonne, G.M., and Christie-Blick, N., 2006, The Ediacaran Period: a new addition to the geologic time scale: Lethaia, v. 39, p. 13–30, doi:10.1080/00241160500409223. Le Heron, P.D., Cox, G., Trundley, A., and Collins, A.S., 2011, Two Cryogenian glacial successions compared: Aspects of the Sturt and Elatina sediment

records of South Australia: Precambrian Research, v. 186, p. 147–168, doi:10.1016/j.precamres.2011.01.014. Link, P.K., 1982, Geology of the Upper Proterozoic Pocatello Formation, Bannock Range, Southeastern Idaho [Ph.D. dissertation]: University of California, Santa Barbara, 131 p. Link, P.K., 1983, Glacial and tectonically influenced sedimentation in the upper Proterozoic Pocatello Formation, southeastern Idaho, in Miller, D.M., et al., eds., Tectonic and stratigraphic studies in the eastern Great Basin: Geological Society of America Memoir 157, p. 165–181. Link, P.K., 1987, The Late Proterozoic Pocatello Formation: A record of continental rifting and glacial marine sedimentation, Portneuf Narrows, southeastern Idaho, in Beus, S.S., ed., Centennial field guide volume 2: Boulder, Colorado, Geological Society of America, p. 139–142. Link, P.K., Miller, J.M.G., and Christie-Blick, N., 1994, Glacial-marine facies in a continental rift environment: Neoproterozoic rocks of the western United States Cordillera in Deynoux, M., Miller, J.M.G., Domack, E.W., Eyles, N., Fairchild, I.J., and Young, G.M., eds., International Geological Correlation Project 260: Earth’s Glacial Record, p. 29–59. Link, P.K., Dehler, C.M., Yonkee, A., and Keeley, J.A., 2011, Systematic regional patterns in detrital-zircon populations from Cryogenian, Ediacaran and Cambrian Sandstones, Brigham Group and tintic quartzite, Northern Utah Thrust Belt: Geological Society of America Abstracts with Programs, v. 43, no. 4, p. 56. Lorentz, N.J., Corsetti, F.A., and Link, P.K., 2004, Seafloor precipitates and C-isotope stratigraphy from the Neoproterozoic Scout Mountain Member of the Pocatello Formation, southeast Idaho: implications for Neoproterozoic earth system behavior: Precambrian Research, v. 130, p. 57–70, doi:10.1016/j.precamres.2003.10.017. Mount, J.F., and Kidder, D., 1993, Combined flow origin of edgewise intraclast conglomerates: Sellick Hill Formation (Lower Cambrian), South Australia: Sedimentology, v. 40, p. 315–329, doi:10.1111/j.1365-3091.1993 .tb01766.x. Myrow, P.M., Tice, L., Archuleta, B., Clark, B., Taylor, J.F., and Ripperdan, R.L., 2004, Flat-pebble conglomerate: its multiple origins and relationship to metre-scale depositional cycles: Sedimentology, v. 51, p. 973–996, doi:10.1111/j.1365-3091.2004.00657.x. Petterson, R., Prave, A.R., Wernicke, B.P., and Fellick, A.E., 2011, The Neoproterozoic Noonday Formation, Death Valley region, California: Geological Society of America Bulletin, doi:10.1130/B30281.1 (in press). Reed, M., Perry, K., Johnston, S.M., and Gehrels, G.E., 2010, Detrital Zircon Provenance of Precambrian–Cambrian Miogeoclinal Sedimentary Rocks, Nevada-Utah Border: Geological Society of America Abstracts with Programs, v. 42, no. 4, p. 66. Rodgers, D.W., Long, S.P., McQuarrie, N., Burgel, W.D., and Hersley, C.F., 2006, Geologic Map of the Inkom Quadrangle, Bannock County, Idaho: Idaho Geologic Survey. Smith, L.H., Kaufman, A.J., Knoll, A.H., and Link, P.K., 1994, Chemostratigraphy of predominantly siliciclastic Neoproterozoic successions: a case study of the Pocatello Formation and Lower Brigham Group, Idaho, USA: Geological Magazine, v. 131, p. 301–314, doi:10.1017/ S0016756800011079. Servais, T., 1996, Some considerations on Acritarch classification: Review of Palaeobotany and Palynology, v. 93, p. 9–22, doi:10.1016/0034 -6667(95)00117-4. Stewart, J.H., 1972, Initial deposits in the Cordilleran Geosyncline: Evidence of a Late Precambrian (<850 m.y.) Continental Separation: Geological Society of America Bulletin, v. 83, p. 1345–1360, doi:10.1130/0016 -7606(1972)83[1345:IDITCG]2.0.CO;2. Stewart, J.H., Gehrels, G.E., Barth, A.P., Link, P.K., Christie-Blick, N., and Wrucke, C.T., 2001, Detrital zircon provenance of Mesoproterozoic to Cambrian arenites in the western United States and Northwestern Mexico: Geological Society of America Bulletin, v. 113, no. 10, p. 1343–1356, doi:10.1130/0016-7606(2001)113<1343:DZPOMT>2.0.CO;2. Vidal, G., and Knoll, A.H., 1983, Proterozoic Plankton: Geological Society of America Memoir 161, p. 265–277. Zhou, C., Tucker, R., Xiao, S., Peng, Z., Xunlai, Y., and Chen, Z., 2004, New constraints on the ages of Neoproterozoic glaciations in south China: Geology, v. 32, p. 437–440, doi:10.1130/G20286.1.

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The Geological Society of America Field Guide 21 2011

Reinterpreted history of latest Pleistocene Lake Bonneville: Geologic setting of threshold failure, Bonneville flood, deltas of the Bear River, and outlets for two Provo shorelines, southeastern Idaho, USA Susanne U. Janecke* Robert Q. Oaks Jr.* Department of Geology, 4505 Old Main Hill, Utah State University, Logan, Utah 84322-4505, USA

ABSTRACT Geologic, geomorphic, and geophysical analyses of landforms, sediments, and geologic structures document the complex history of pluvial Lake Bonneville in northern Cache Valley, NE Great Basin, and shows that the outlet of Lake Bonneville shifted ~20 km south after the Bonneville flood. The Riverdale normal fault offsets Bonneville deposits, but not younger Provo deposits ~25 km southeast of Zenda, Idaho. Rapid changes in water level may have induced slip on the Riverdale fault shortly before, during, or after the Bonneville flood. Although other processes may have played a role, seismicity might have been the main cause of the Bonneville flood. The outlet of Lake Bonneville shifted south from Zenda first 11, then another 12 km, during the Provo occupation. The subsequent Holocene establishment of the drainage divide at Red Rock Pass, south of Zenda, resulted from an alluvial fan damming the north-sloping valley. Weak Neogene sediments formed sills for the three overflowing stages of the lake, including the pre-flood highstand. Field trip stops on flood-modified landslide deposits overlook two outflow channels, examine and discuss the conglomerate-bearing sedimentary deposits that formed the dam of Lake Bonneville, sapping-related landforms, and the Holocene alluvial fan that produced the modern drainage divide at Red Rock Pass. The flood scoured ~25 km of Cache and Marsh Valleys, initiated modest-sized landslides, and cut a channel north of a new sill near Swan Lake. Lake Bonneville dropped ~100 m and stablilized south of this sill at the main, higher ~4775 ± 10 ft (1456 ± 3 m) Provo shoreline. Later Lake Bonneville briefly stabilized at a lower ~4745 ± 10 ft (1447 ± 3 m) Provo sill, near Clifton, Idaho, 12 km farther south. An abandoned meandering riverbed in Round Valley, Idaho, shows major flow of the large Bonneville River northward from the Clifton sill. Field trip stops at both sills and overlooking the meander belt examine some of the field evidence for these shorelines and sills.

*[email protected]; [email protected]. Janecke, S.U., and Oaks, R.Q., Jr., 2011, Reinterpreted history of latest Pleistocene Lake Bonneville: Geologic setting of threshold failure, Bonneville flood, deltas of the Bear River, and outlets for two Provo shorelines, southeastern Idaho, USA, in Lee, J., and Evans, J.P., eds., Geologic Field Trips to the Basin and Range, Rocky Mountains, Snake River Plain, and Terranes of the U.S. Cordillera: Geological Society of America Field Guide 21, p. 195–222, doi:10.1130/2011.0021(09). For permission to copy, contact [email protected]. ©2011 The Geological Society of America. All rights reserved.

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Janecke and Oaks The Bear River, which enters Cache Valley at the mouth of Oneida Narrows, 17 km ENE of the Clifton sill, was the main source of water in Lake Bonneville. It produced 3 sets of deltas in Cache Valley—a major delta during the Bonneville highstand, a larger composite delta during occupation of two Provo shorelines, and at least one smaller delta during recession from the Provo shoreline. The Bonneville delta and most of the Provo delta of the Bear River were subaqueous in Cache Valley, based on their topsets being lower than the coeval shorelines. The Bonneville delta is deeply dissected by closely spaced gullies that formed immediately after the Bonneville flood. The delta morphologies change sequentially from river-dominated to wave-dominated, then back to river-dominated. These unique shapes and the brief, intense erosion of the Bonneville delta record temporal changes in wave energy, erosion, vegetation, and/or storminess, at the end of the Pleistocene. Stops on a delta near Weston, Idaho, reveal many of the distinguishing features of the much larger deltas of the Bear River in a smaller, more concentrated form. We will see and discuss the ubiquitous gully erosion in Bonneville landforms, the nearly undissected Provo delta, the subaqueous topset of the Provo delta, and the wave-cut and wave-built benches and notches at the upper and lower Provo shorelines.

INTRODUCTION AND SETTING The great Bonneville megaflood in the northeast Great Basin, 17,400 calendar years ago, released roughly half the water from this deep, large pluvial lake (Gilbert, 1880, 1890; Malde 1968; Currey, 1990; Oviatt et al., 1992; O’Connor, 1993; Hart et al., 2004; calibration to calendar years in part from Guido et al., 2007). Approximately 4750 km2 of water drained northward into the Snake River with a maximum discharge of about a million cubic meters per second (Figs. 1, 2, 3, and 4; Malde, 1968; O’Connor, 1993). Massive scouring and redeposition are well documented north of the outlet (Gilbert, 1880, 1890; Malde, 1968; O’Connor, 1993). After the flood, the Provo shoreline in the Bonneville basin was established ~102 m (335 ft) lower than the highstand Bonneville shoreline (Fig. 3). Gilbert (1880, 1890) identified the threshold at the time of the Bonneville flood at the site of Zenda, Idaho. Gilbert (1890, p. 178) concluded that the outlet of Lake Bonneville at the Provo shoreline shifted south at least 11 km to a position “…between Swan Lake and the Round Valley Marsh” after the Bonneville flood (Figs. 1, 2, 4, 5). Subsequent workers concluded instead that the outlet shifted only 2.5 km south, from Zenda to the modern divide at Red Rock Pass (Ives, 1948; Williams, 1962a, 1962b; Williams and Milligan, 1968; Burr and Currey, 1988; Smith et al., 1989; Bright and Ore, 1987; O’Connor, 1993; Link et al., 1999). We argue here that Gilbert (1880, 1890) was correct. In addition, a second outlet southeast of Clifton, Idaho, is required to explain relationships in Round Valley, 12 km south of Swan Lake (Figs. 7 and 9). Surprisingly little is known about the geologic record of the Bonneville flood and its aftermath near its threshold in northern Cache Valley and southern Marsh Valley (Figs. 1–6, 8, 10). The cause of the catastrophic failure in southernmost Marsh Valley is still uncertain despite >120 years of study following the excep-

tional work of G.K. Gilbert (1880, 1890). Hypotheses that have been advanced for the flood include the failure of an alluvial sill, sapping, failure of porous karst, involvement of Paleozoic bedrock, and landsliding (Gilbert, 1890; Ives, 1948; O’Connor, 1993; Link et al., 1999; J. Stewart Williams, 1967, oral commun. to Oaks; Link et al., 1999; Malde, 1968; Cornwell and Shroder, 1990; O’Connor, 1993). Our work suggests seismicity as an additional potential triggering mechanism. Reports of shoreline altitudes in northern Cache Valley are inconsistent, with the greatest discrepancy in the altitude of the Provo shoreline (Hardy, 1957; Bright, 1966; Link, 1982a, 1982b; Burr and Currey, 1988; Currey and Burr, 1988; Smith et al., 1989; Godsey et al., 2005; Williams and Milligan, 1968). Early mappers identified a shoreline between 4775 and 4800 ft as the Provo shoreline, whereas later workers favored a lower altitude between 4737 and 4756 ft (Hardy, 1957; Bright, 1966; Williams and Milligan, 1968; Link, 1982a, 1982b; Burr and Currey, 1988; Currey and Burr, 1988; Smith et al., 1989; Godsey et al., 2005). Rebound complicates the correlation of shorelines in many parts of the Bonneville basin, but there was little rebound in northernmost Cache Valley because it was at the margin of the lake and shallow (Crittenden, 1963; Bills et al., 1994). The lower shoreline altitude of 4737–4756 ft for the Provo shoreline has a closure far south of the divide at Red Rock Pass, where most workers have interpreted the threshold after the Bonneville flood (Fig. 2; Currey et al., 1984). The deltas of the Bear River, nearby, have not been mapped nor interpreted in a landscape context, and stratigraphic analyses of the Bonneville delta have produced varied results (Figs. 1, 2, 11, and 12) (Oriel and Platt, 1980; Lemons et al., 1996; Lemons, 1997; Anderson, 1998; Anderson and Link, 1998; Milligan and Lemons, 1998; Milligan and Chan, 1998; Lemons and Chan, 1999). All of these factors, plus our chance discovery of a large relict meander belt 15–25 km south of the original

Reinterpreted history of latest Pleistocene Lake Bonneville

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Figure 1. Map of lake levels within Cache Valley. Deltas of the Bear River are differentiated from other lacustrine deposits. Provo shoreline is simplified west of Cache Valley. RNF—Riverdale normal fault (landslide?). Major structures and locations of other maps are shown.

10 km

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Fig. 10 Marsh Valley Fig.. 8

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Between highest Bonneville shoreline and upper Provo shoreline Between upper and lower Provo shorelines Below lower Provo shoreline subProvo deposits

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Figure 2. Digital elevation model delineating the main features of northern Cache Valley. Color schemes are keyed to elevations of the Provo and Bonneville shorelines. Colors are same in the following maps except for minor adjustments to account for rebound. Notice the part of the Swan Lake scour channel that has been filled and dammed by the Holocene fan of Marsh Creek near Red Rock Pass. The tan colors within the Swan Lake scour channel reflect the filling of this area above the altitude of the 4775 ft (1456 m) Provo shoreline.

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Bear River

Fig.. 12 Provo delta top

Bear River delta

Provo delta top

Weston delta Cub River

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Cache Valley

Reinterpreted history of latest Pleistocene Lake Bonneville threshold of Lake Bonneville prompted our detailed reevaluation of this fascinating body of water and its erosional and depositional history. Much of our work on this topic is presented in a manuscript that is currently in the review process. In that paper, we provide detailed evidence for the two sills, two important shorelines at the Provo level, two discharge channels north of Red Rock Pass, modest-sized landslides, and, most importantly, the possibility of the Bonneville flood resulting from an earthquake on the Riverdale fault zone. In order to limit redundancy, those data and interpretations are summarized here in an abbreviated form, and the evidence for most of these interpretations will be discussed at

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field trip stops. We present additional data and analysis of the deltas of the Bear River and their relevance for interpreting changing climate in the latest Pleistocene. Methods We examined landforms of northern Cache Valley on aerial photographs, orthophoto-quadrangles, digital elevation models, and topographic maps. Hundreds of topographic profiles and 10-m digital elevation models created in GeoMapApp were analyzed (Figs. 2, 5, 10, 11, and 12). Geologic mapping (Mayer, 1979; Oriel and Platt, 1980; Link, 1982a, 1982b; Link and LeFebre,

Figure 3. Hydrograph of Lake Bonneville modified to incorporate results from this study. Original data from Oviatt et al. (1992) and Godsey et al. (2005) were calibrated to calendar years based on Marrero (2009) and Guido et al. (2007). A drop of lake level during Provo time (Godsey et al., 2005) might have separated the 4775 ft (1456 m) and 4745 ft (1445 m) Provo shorelines, as shown, but this is uncertain. Isotopic data from caves in Arizona (lower curves) (Wagner et al., 2010) provide evidence for glacial conditions during the Bonneville highstand and subsequent Bonneville flood. The abrupt warming of climate during the Provo occupation might have contributed to incision of the Swan Lake sill by increased runoff from melting glaciers in the Uinta Mountains.

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We t scour Wes West scou sc cou our channel chan hannel ne e (40-50 (4 40-5 -5 50 m lower) lo lower ower w )

Red Rock Butte

Figure 4. Photograph looking north-northeast at the lower, partly infilled western scour channel (left), west of Red Rock Butte, and the higher eastern scour channel (right). Located near Stop 3.

Janecke and Oaks

East point 40-50 higher) E t sscour Eas cour channel (low poi o ntt is 40 4050 m high 50 h igh igher gher) er) er

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1983; DeVecchio, 2002; Hennings et al., 2001; DeVecchio et al., 2003; Carney et al., 2003; Carney and Janecke, 2005; Steely and Janecke, 2005; Steely et al., 2005; Long and Link, 2007; Janecke and Oaks, unpublished mapping) and analysis of topography and of gravity data (Peterson and Oriel, 1970; Stanley, 1972; Scheu, 1985a, 1985b; Kruger et al., 2003; Eversaul, 2004; Oaks et al., 2005) provided constraints including the subsurface positions of bedrock-cored highs in northern Cache Valley. We report the present range of shoreline altitudes of both erosional and depositional features in the study, uncorrected for rebound, because there has been little differential rebound of the Provo shoreline here (Solomon, 1999). We remapped key areas near Red Rock Pass (Fig. 13) and in selected localities elsewhere. Oaks examined hundreds of drillers’ logs of water wells in northern Cache Valley for the grain sizes of the sediments and depths to contacts with Neogene Salt Lake Formation and to Paleozoic or Proterozoic bedrock, and constructed geologic cross sections from them. Soil maps helped us to identify some features. RESULTS Shorelines The highest Bonneville shoreline is mostly erosional around northern Cache Valley, as wave-cut scarps and terraces, locally with distal wave-built terraces (e.g., Gilbert, 1890) at the bases of small triangular facets. It lies between 5090 and 5120 ft (1551– 1561 m) and is very prominent, except in northernmost Cache Valley and southernmost Marsh Valley (Bright, 1963). After the Bonneville flood, the lake stabilized at a shoreline between 4760 and 4780 ft (1451–1457 m) in northern Cache Valley, the main Provo shoreline of Lake Bonneville. We refer to this higher, older stand as the “4775 ft Provo shoreline” (1456 m) (Figs. 3 and 7). A younger, lower Provo shoreline, between 4740 and 4750 ft (~1447 m), is much more subtle than the higher Provo shoreline around northern Cache Valley (Figs. 3 and 9), and in many areas does not form a clear or continuous shoreline. However, the lower Provo shoreline is well expressed on the Weston delta as a step (~9 m) in the smooth topset of the Provo delta (Fig. 12; Stop 8). The main higher Provo shoreline is 30 ft higher than inferred in most publications after 1970. North of Red Rock Pass The divide at Red Rock Pass (Figs. 4, 5, 6, and 8; Stop 3), between Marsh Creek and Cache Valley, separates the Great Basin from the Columbia River drainage basin. The divide was located close to Zenda prior to the Bonneville flood (Gilbert, 1890). It was thought that there had been outflow during the highstand of Lake Bonneville prior to the discovery of groundwaterrelated landforms in Marsh Valley in the Marsh Creek alluvial fan (O’Connor, 1993; we show below that this is mostly a pediment with fairly thin alluvial and eolian cover). That discovery raised the possibility that Lake Bonneville was closed until sapping

Reinterpreted history of latest Pleistocene Lake Bonneville weakened an alluvial dam near Zenda and initiated outflow and the Bonneville flood (O’Connor, 1993). We verified and mapped the sapping-related landforms in southern Marsh Valley, and also reconstructed the landscape by projecting the regular pediment and alluvial-fan surfaces related to Aspen Creek (west of Marsh Creek) and Marsh Creek (east and north of Marsh Creek), across the fairly narrow channel cut by the Bonneville flood. Sapping modified the surface of the Marsh Creek pediment and alluvial fan from the altitude of the Bonneville shoreline down to the floor of Marsh Valley (Fig. 6). The reconstructed landforms show that Lake Bonneville probably had an outlet when it occupied its highest shoreline because the high pediments and alluvial fans coalesced at the altitude of the Bonneville shoreline, not above it. The co-occurrence of sapping-related landforms and overland flow is enigmatic. Perhaps it reflects a short period of subsurface throughflow before overland flow was established at Zenda. Possibly the subsurface throughflow reduced the overflow to an intermittent occurrence during high-water years, and thus prolonged the highstand near that level prior to final collapse. Possibly the graded, northwestsloping pediment surface allowed moderate overflow without downcutting, similar to the Madras canals of India that neither erode nor deposit sediment (Leopold et al., 1964).

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A modest volume of landslide material is preserved near the scoured channels in the Red Rock Pass area, but the volume was much less than has been suggested previously (Sewell, cited in Smith et al., 1989). There was an early Bonneville discharge channel at Red Rock Pass, east of Red Rock Butte, that helped to empty the Bonneville basin of water before and during the flood. It was deactivated before the end of the flood. Our new mapping shows that all but the uppermost few tens of meters of material in the area of the Zenda dam are cemented conglomerate and interbedded silt and sand of the Neogene Salt Lake Formation (Stop 6). Paleozoic rocks do not crop out at the pre-Bonneville divide, so models of the Bonneville flood that involve flow of water through karst in Paleozoic carbonate (J. Stewart Williams, 1967, oral commun. to Oaks; Link et al., 1999) cannot be correct. Quaternary alluvial sediments overlying the pediments thicken and thin in cut-and-fill geometries, and thicken valleyward near Marsh and Aspen Creeks, but do not appear to have comprised much of the Zenda threshold. Results from South of Red Rock Pass The northern third of Cache Valley is complex and variable in its geomorphology and bedrock structure (Figs. 1, 2, 5,

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Figure 5. Detailed topographic map and profile showing the two discharge channels that were produced by the Bonneville flood north of Red Rock Pass. The eastern channel stopped incising after ~60 m of downcutting from the Bonneville highstand, whereas the western channel experienced the full ~100 m of incision and some younger infilling. Colors match Figs. 2 and 3.

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Figure 6. Key features of the Bonneville highstand and the Bonneville flood, overlain on a geologic map of southern Marsh Valley, with the divide between Marsh and Cache Valleys and the Bonneville highstand prior to the flood (modified from DeVecchio et al., 2003). Stops 3−6 are in this area.

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202 Janecke and Oaks

Reinterpreted history of latest Pleistocene Lake Bonneville

203

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Figure 7. Color aerial photograph from Google Earth shows locations of Stops 1, 2, 7, and 8, and their relationships to the Swan Lake horst, the Twin Lakes horst block, Round Valley and its relict meander belt, and sills at Swan Lake and Clifton. Note that Swan Lake is a residual low in the Swan Lake scour channel that lies between younger infilling alluvial fans. Blue outlines define the east and west edges of the relict meander belt. White lines denote the east and west edges of scoured terrain, yellow shows the approximate outline of Round Valley, and the white dashed line delineates the highest parts of the Twin Lakes horst. The latter continues east and west in the shallow subsurface according to gravity data (Eversaul, 2004).

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6, and 7). We identify five distinct geomorphic zones there, from: (1) the open, former lake bottom of Lake Bonneville south of Preston, (2) the large deltas of the Bear River which have a southern margin 3 km south of Preston and persist north-northwest 12 km to a constriction near Clifton, (3) the Twin Lakes horst block and associated low hills, (4) Round Valley, and (5) the hills and valleys of the Swan Lake terrain in northernmost Cache Valley. Scouring from the Bonneville flood is confined to the Swan Lake terrain and southern Marsh Valley. Shorelines are best developed south of the deltas of the Bear River (see Stop 9). The higher Provo shoreline is mappable as far north as Swan Lake, Idaho (Fig. 2; Gilbert,

sapping Stop 5

Round Valley Terrain, Meander Belt, and Riverdale Normal Fault

Aspen bench pediment

Stop 4

Marsh Creek

no sapping

Stop 6

1890; Bright, 1963; Stop 2), whereas the lower Provo shoreline is only expressed south of the Twin Lakes horst (Curry et al., 1984). The Twin Lakes hills were an irregular island in Lake Bonneville during the Bonneville highstand and the subsequent higher Provo level (Figs. 2 and 7). Round Valley preserves a marshy low area that once contained a large north-flowing river. The description of the relationships in Round Valley is included here because they require a major re-interpretation of Lake Bonneville. Pre-Cenozoic bedrock underlies the Twin Lakes horst and the ridges in the Swan Lake terrain. Gravity data (Eversaul, 2004) show where the bedrock is close to the surface in areas of thick surficial cover (see Stops 1–9 for more detail). The Salt Lake Formation is the basin fill that was coeval with slip on the Bannock detachment fault and it makes up the sills of Lake Bonneville (Link, 1982a, 1982b; Janecke and Evans, 1999; Oaks et al., 1999; Oaks, 2000; Janecke et al., 2003; DeVecchio, 2002; Carney and Janecke, 2005; Steely and Janecke, 2005; Steely et al., 2005; Oaks).

Stop 3

Bannock Range

Alluvial fan, Holocene Figure 8. Locations of Stops 3, 4, 5, and 6 and their relationships to the Marsh Creek pediment and alluvial fan, the east and west scour channels, highest sapping-related landforms, the maximum extent of flood-related scouring (white lines), and young alluvial fans (brown) that partly filled the now-dry bed of the Bonneville River. The current drainage divide at Red Rock Junction is the result of differential infilling of the scoured area by Marsh Creek (Gilbert, 1890).

Round Valley is bounded by bedrock ridges and a mountain range on three sides, and abuts the foreset of the Bonneville delta of the Bear River on its east side. Its floor is nearly flat, and there is a break in slope around its perimeter that is near the altitude of the higher 4775 ft (1456 m) Provo shoreline. The center of the valley is as low as 4740 ft (1444 m). Most of Round Valley lies between 4745 and 4750 ft (~1447 m) altitude. Round Valley preserves landforms of an ancient meandering river (Figs. 7 and 9; see Google Earth; Stop 7). The abandoned meander belt in Round Valley has six meander loops, two to three dozen depositional packets of large curving ridges that resemble point bars that were added northward, oxbow lakes, and several tributaries that enter from south to north. The meander belt connects directly into the Swan Lake scour channel at its northeast end (Figs. 2 and 7; near Stop 2). The channel and floodplain of this large river system formed below, and after, the higher Provo shoreline. The wavelengths of the meanders are ~2.0–2.5 km, and average 2.3 km (7 measurements from Fig. 9), with an average radius of ~0.6 km (Fig. 5). The floodplain produced by lateral accretion is up to ~2.75 km wide, and there are 5–8 accretionary point-bar complexes visible in each meander loop (Fig. 5). The width of the relict channel of the meander belt in Round Valley is ~140 ft (43 m) (average of 12 measurements along the channel courses). The average meander wavelength and average channel width permit other paleohydraulic parameters to be calculated: (1) channel depth ~1.5 m (Dury, 1965); (2) channel slope ~0.1 m/km (Schumm, 1972); (3) mean annual discharge ~6 m3/sec (Schumm, 1972); and (4) mean annual flood ~25 m3/sec (Schumm, 1972). These calculations lie within the data sets shown in Leopold et al. (1964), and the ratio of slope to discharge is within their field for meandering streams rather than braided to straight streams.

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Figure 9. Color aerial photograph from Google Earth of the relict meander belt in Round Valley.

Reinterpreted history of latest Pleistocene Lake Bonneville 205

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The Riverdale normal fault is buried beneath the northeast corner of Round Valley, where a vegetation lineament overlies its probable trace. It separates the more open part of Cache Valley, including the relict meanders in Round Valley, in its hanging wall, from the linear and narrow Swan Lake scour channel in its footwall. To the southeast, it lies near the top of a prominent SWsloping gravity gradient along the NE edge of the relatively flat floor of Cache Valley (Figs. 1, 2, and 11; Eversaul, 2004; Oaks et al., 2005). Linear gullies are eroded along its trace across the

foresets of the highstand Bonneville delta north of the Bear River (see Stop 1). The Riverdale fault lies in the foreset of the Bonneville delta and has a position reminiscent of delta-front landslides (Paolo, 2000; Heller et al., 2001; Paola et al., 2001). We cannot rule out a landslide interpretation but a fault model explains more of the known relationships. The Clifton sill is at the south end of the meander belt in Round Valley. It coincides with the boundary between undissected fluvial landforms, to the north, and a smooth Provo top-

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Figure 10. Topographic map of the deeply dissected landscape north of the Swan Lake horst. The similar width and depth of the Swan Lake scour channel along this entire reach supports Gilbert’s interpretation of a Provo outlet of Lake Bonneville located between Swan Lake and Round Valley. Note that the only part of this channel that is above the Provo altitude is near Red Rock Pass, where a Holocene alluvial fan has raised the floor of the channel.

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Figure 11. Topographic maps of the deeply dissected Bonneville delta of the Bear River (orange), little dissected topset of the Provo delta of the Bear River (greens) and the Riverdale fault zone cutting across foresets of the Bonneville delta (between arrows). Notice that the Provo deposits, shown in green and light blue, are not cut by the Riverdale fault along strike to the north. The profile shows the deep and regular gullies within the Bonneville delta and their complete absence in the Provo delta. Contour interval is 10 m.

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Reinterpreted history of latest Pleistocene Lake Bonneville 207

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Figure 12. Top: Southern margin of the Provo delta of the Bear River is highlighted to show its topset (orange), upper foreset (yellow and blue), and lower foreset and bottomset (dark blue). Note the concave south shape of the delta front and the few streams that have incised this landform since its abandonment in the latest Pleistocene. At least one younger delta was inset onto the delta front of the large composite Provo delta during the final recession of Lake Bonneville, and the erosional remnants complicate the relationships near the incised Bear River. Bottom: Enlarged view of the three key shorelines of Lake Bonneville in the Weston, Idaho, area. Note the main higher Provo shoreline is higher than the topset of the delta at Weston and bounds the base of strongly gullied hillslope of Salt Lake Formation (Tsl). Gullies like this are very common between the Bonneville and higher Provo shorelines throughout the region.

Figure 13. Possible reconstruction of events in the southern part of Marsh Valley and Red Rock Pass area. Bolded events are not depicted here because they were south of this area. Stage 1: Overflow of Lake Bonneville stabilized the position of the lake, and produced a prominent highest Bonneville shoreline. The Bonneville River of Gilbert (1890) was born at this time with an outlet near 1550 m. An earthquake on the Riverdale fault might have started the Bonneville flood, possibly with seiche waves. Stage 2: Early during the Bonneville flood the Bonneville River in the eastern scour channel cut down ~60 m through the Zenda threshold. Salt Lake Formation was exposed in the floor of the channel. Stage 3: A landslide from the east blocked the east scour channel, and flow shifted entirely to a channel west of Red Rock Butte. Stage 4: Recurring landslides into the west channel were washed away as the flood continued. Eventually the flood ended, and a new outlet for the Bonneville River emerged at the Swan Lake area 11 km farther south. Northward flow in the Bonneville River continued within the western channel. Climate change at the end of the glaciation caused a regression from the Provo shorelines. Stage 5. Immediately after the end of the Provo occupation, the Bonneville River ceased to exist. Its dry river bed (shown in purple here even though it is now dry) had become the new base level for the small tributary streams in the area. Now sediment delivered onto the bed of the dry river by those tributaries was able to accumulate, rather than being carried north by the mighty Bonneville River. Through time the Holocene alluvial fans built up enough topography to block the northward gradient of the dry river bed. Marshes and small lakes, like Swan Lake, formed between the Holocene fans. The largest fan was at the confluence of Marsh Creek and the scour channel because Marsh Creek is the largest stream. Its crest became the modern drainage divide at Red Rock Pass.

2.

4.

1.

3.

5.

Contour on pediment surface

Fault scarp and possible fault scarps that may have been modified by flood water

Highest level of sapping features on Marsh Creek fan-pediment

Highest Bonnieville shoreline

Landslide deposit sculpted by the flood Scoured landform parallel to flow

Landslide deposits, reconstructed

Pre-Tertiary bed rock

Pleistocene fan and pediment deposits

East scour channel West Swan Lake scour channel

Threshold near Zenda, Idaho (~18 ka) Drainage Divide at Red Rock Pass (now) Post-Provo alluvial fan

Reinterpreted history of latest Pleistocene Lake Bonneville 209

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set incised by a few deep, steep-sided gullies, to the south (Figs. 1, 2, and 7). Deltas of the Bear River The topsets and foresets/bottomsets (delta front) of the highstand Bonneville delta and the composite Provo-level delta of the Bear River form the largest geographic terrain of northern Cache Valley. The foreset-topset contact, which is typically assumed to coincide with shorelines in Gilberttype sand-and-gravel deltas, is below both the Bonneville and Provo shorelines in these mostly clay-and-silt deltas. There is only one very large delta complex associated with the Provo shorelines, and there is no clear distinction between landforms formed during occupation of the 4775 ft (1456 m) and 4745 ft (1447 m) Provo shorelines (Fig. 3). We therefore treat this composite landform as a single delta, but further study is needed. Sub-Provo deposits of at least two prior lake cycles, plus transgressive deposits from the rise to the Bonneville highstand could be buried in the subsurface of northern Cache Valley beneath the Provo-level deltas. If so, these deposits reduce the amount of sediment needed to fill in the area beneath the Provo and Bonneville highstand deltas. None of the prior lake cycles significantly exceeded the altitude of the Provo shoreline (McCoy, 1987, as modified in Hart et al., 2004), so there is no ambiguity about the age of the highest deltaic deposits. Bonneville Delta of the Bear River and the Riverdale Fault The Bear River reaches Cache Valley through a confined bedrock valley ~16 km long at Oneida Narrows, 27.5 km eastsoutheast of Red Rock Pass (Fig. 1). The highstand Bonneville delta of the Bear River is confined along the northeast edge of Cache Valley and is asymmetric, with more of its surface area north of its inlet where two small creeks contributed sediment to a coalesced delta complex (Figs. 1, 2, and 11). The delta is large, spanning roughly 10 km from its distal point to its inlet into Cache Valley at Oneida Narrows, and 20 km from northwest to southeast. Most of the surface area of the Bonneville delta of the Bear River is in Cache Valley, but a long narrow finger of the delta was first deposited in Gentile Valley and SSW through Oneida Narrows (Fig. 1). Most of that sediment was eroded during and after regression from the Bonneville highstand. Although part of this accumulation may have been swept across the Swan Lake and Clifton sills into the Columbia River Basin, much was probably redeposited into the Provo delta. The highstand Bonneville delta front was originally quite smooth and weakly convex to the southwest. The delta filled in an irregular reentrant in northeast Cache Valley, and then built out into the more open part of the valley, filling it almost halfway from east to west. Relict erosional gullies dissect the originally smooth delta foreset and the distal parts of the topset of the Bonneville delta, both north and south of the presently inset Bear River. The topset

is less dissected by the gullies. Gullies are grossly perpendicular to the delta front, except for those that parallel the Riverdale fault, but there is a considerable range of trends in the gullies (Fig. 11). Spacing, widths, depths, and lengths of gullies are very regular. Gullies that head in topsets of the Bonneville delta do not extend into canyons upgradient (Fig. 11). There is little continuity of the gullies upslope and downslope across the Riverdale fault (Fig. 11). Gullies upslope of that fault trend east-northeast, whereas those downslope of the fault vary, but most have east- to east-southeast trends. Gullies on the highstand Bonneville delta all grade westward to, and end at, the shoreline and adjacent topset of the composite Provo delta of the Bear River. Alluvial fans deposited at the bases of the lower gullies are small, insufficient to contain all of the sediment eroded from those gullies. Most of the material eroded from the gullies must have been reworked and incorporated into the composite Provo delta. The Riverdale fault zone, which is demarked by the gullies that trend across the Bonneville delta, does not cut any of the Provo-age deposits along strike N to S. Thus, the gullies in the Bonneville delta of the Bear River that follow the Riverdale fault are coeval with the Bonneville flood. Provo Delta of the Bear River Front, Foreset, and Bottomset of Provo Delta The unconfined southern front, at the foreset-topset contact of the lower Provo shoreline, is ~16 km wide, east to west across Cache Valley (Figs. 1, 2, and 12). The change from the gently sloping topset and the steeper delta front (foreset) lies between 1430 and 1439 m (Figs. 2 and 12). This foreset has a very gentle slope southward. The base of the delta foreset coincides with a decrease in slope and a change from slightly more sand-rich sediment of the delta to more silt- and clay-rich sediment of the delta’s bottomset (Web Soil Survey, accessed 30 June 2008; http://websoilsurvey.nrcs.usda.gov/app/WebSoilSurvey.aspx). The southern front of the Provo delta has a smooth, open, concave-south shape (Figs. 1, 2, and 12) that is typical of a wavedominated delta. A small bend southward in the middle of the delta front at lower altitudes likely resulted from outbuilding of at least one younger delta as the lake fell below the Provo level. The fine sand, silt, and clay deposited by the Bear River in N Cache Valley was easily redistributed by waves and currents, unlike the heavier gravel that dominates deltas at the mouths of steep mountain streams farther south in Cache Valley (Gilbert, 1890). Topset of the Provo Delta The topset of the higher Provo delta is obscured by fluvial deposits in Round Valley and is interrupted by irregular topography of the Twin Lakes horst (Figs. 1, 2, 11, and 12). Logs of water wells, which penetrate the underlying Salt Lake Formation in some cases, show that most of the sediment in the Provo delta complex is clay, silt, and fine sand. The entire northeast and western edges of the Provo delta complex lap against the older Bonneville delta or bedrock composed of Salt Lake Formation,

Reinterpreted history of latest Pleistocene Lake Bonneville Paleozoic or Precambrian rocks, respectively, and the only unconstrained depositional margin is its southern east-trending delta front (Figs. 1, 2, and 12). Deposition was probably rapid at first as voluminous unconsolidated sediment upstream from Provo shoreline was remobilized. Some of the topset of the Provo delta west of the Twin Lakes horst has faint relict meander scars. Most of the rest of the topset is smooth. Most of it likely formed underwater because the Provo shorelines are higher in altitude around its perimeter (Fig. 1, 2, 11, and 12; Currey et al., 1984). Dune-Covered Subdelta(s) below the Provo Delta One or two small, lobate, post-Provo deltas were formed by the incising Bear River on the southern edge of the Provo delta during recession below the Provo shoreline (Figs. 1, 2, and 12). Further incision removed large parts of these lobate deltas. Erosional remnants of these deltas are mostly on the east side of the Bear River. Together the post-Provo delta remnants cover ~40 km2 southwest of Preston, with the altitude of the higher, northern topset at ~4660 ft (~1420 m) and the high point on the lower, southeastern delta at ~4582 ft (~1397 m). There are younger parabolic sand dunes, formed by winds from the southwest, that obscure much of the post-Provo delta surfaces and some of the distal topset of the Provo delta southwest of Preston. The post-Provo delta is lobate in plan view, and its convexsouth margin differs from the smoother cuspate to gently curving margins of the Provo and Bonneville deltas of the Bear River (Figs. 1, 2, 11, and 12). The dune-covered subdeltas are composed of fine sediment derived from erosion of the adjacent Provo delta, but exposures show significant sand. The post-Provo delta may contain a greater proportion of sand due to winnowing by the Bear River confined within an inset valley. Thus, its lobate, southward-convex shape might be due to an overall increase in grain size plus concentration of deposition in a small area. DISCUSSION What Caused the Bonneville Flood: Overland Flow, Sapping, an Earthquake, or a Landslide? It is challenging to pin down the main trigger for the Bonneville flood and to develop a definitive test because many processes could have triggered it, and most of the evidence is eroded near the site of failure. Sapping-related dam collapse provides one plausible trigger of the Bonneville flood (O’Connor, 1993). However, incision by overland flow just as easily could have triggered the Bonneville flood, and overland flow during the Bonneville highstand is likely. The cross-cutting relationship between the Riverdale normal fault and the highstand Bonneville foreset deposits of the Bear River raises another intriguing possibility, that a moderate to major earthquake on that fault (or emplacement of a lateral spread with a headscarp at the trace of the Riverdale fault) triggered the flood. After all, the Riverdale fault (landslide?) cuts and deforms deltaic deposits and landforms that

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predate the flood, yet does not deform any deposits or landforms that postdate the flood. An earthquake on the Riverdale (or other nearby) fault could have produced a seiche wave that overtopped the Zenda dam with high-velocity waters of sufficient volume to destabilize that dam, liquefy the dam, rapidly incise the length of the threshold, or cause other critical damage. Theoretical considerations suggest that the rapid drawdown of Lake Bonneville could have been a “perfect storm” for triggering reservoir-induced seismicity. Numerous small (and some large) earthquakes develop above the background frequency as bodies of water fill and empty, notably at dams (Gupta, 2002; Telesca, 2010). For this reason, and because the cross-cutting relationships allow it, it also possible that the opposite sequence of events occurred, i.e., that the Bonneville flood unloaded the crust rapidly, and thereby induced slip on the Riverdale fault (landslide?). As ~100 m of water was abruptly removed from the hanging wall of the fault zone, an earthquake may have initiated in response to residual elevated pore pressures and lower vertical loads. Either way, differential loading and unloading is likely to have induced seismicity throughout the Bonneville Basin when Lake Bonneville was rising to its highest levels and during the rapid recession afterwards (Fig. 3). Evidence of this process is documented in the Salt Lake City and Brigham City segments of the Wasatch fault, which ruptured during the highstand of Lake Bonneville (McCalpin, 2002; McCalpin and Forman, 2002). Some large landslides immediately west of Cache Valley also seem to coincide with this time period (Biek et al., 2003). Trenching and dating of the Riverdale structure are required to test these competing models. There are other possible triggers for failure of the Zenda threshold that are less likely to have been the main trigger for the flood. Possible Effect of Climate Change on the Morphology of Deltas of the Bear River The contrast of intense and widespread gullying of the highstand Bonneville delta and little or no erosion of the Provo delta complex suggests that the local climate abruptly changed shortly after the Bonneville flood ca. 17.4 cal yr BCE. This change inhibited widespread gullying during the latest Pleistocene and transition into the Holocene. Prior analyses of the climate in Utah during the last glacial epoch suggested a 15%–30% increase in precipitation relative to the present, so that much of the growth of Lake Bonneville probably was due to increased cloudiness and reduced evaporation in lower air temperatures (Lemons et al., 1996; Milligan and Chan, 1998). Although the total amount of precipitation in the area of Lake Bonneville may not have increased greatly, the snowline in the headwaters of the Bear River in the Uinta Mountains lowered considerably (Munroe et al., 2006). Thus, there was probably an increase in runoff and more sediment (including much fine-grained glacial flour) carried by the Bear River during glacial times. Furthermore, more of the total precipitation probably accumulated as snow. With lower evaporation and more snow

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pack, the frequency and intensity of floods in the springtime should have been greater than at present. Perhaps dewatering and piping-related gully formation explains the ubiquitous gullies in sediment between the Bonneville and Provo level (Joel Pederson, 2007, oral commun.). If dewatering were important, we would expect to observe theatershaped and scalloped heads and walls along the gullies that dissect the Bonneville deposits instead of the very regular linear, parallel V-shaped gully systems that formed there. In addition, gully development by dewatering should not be very pronounced in rather impermeable materials such as the tuffaceous parts of the Salt Lake Formation. Hillslopes underlain by tuffaceous Salt Lake Formation have many V-shaped gullies between the Bonneville and Provo shorelines (Fig. 12). Overall, four processes may have resulted in more erosion of the Bonneville delta than the Provo delta: (1) more rapid exposure of erodable fine sediment during the Bonneville flood, and the inability of plants to rapidly colonize and stabilize the newly exposed hillslopes after the Bonneville flood due to continued cold climate; (2) slightly more precipitation, primarily as snow, during the glacial (and lake) maximum; (3) more intense storms when the continental glaciers were nearby in present northern Idaho and Montana; and (4) more frequent, more voluminous, and more erosive springtime flooding. Some combination of these processes was probably operating. Wind Analysis of spits and shorelines built into Lake Bonneville typically indicates strong north winds (Schofield et al., 2004; Jewell, 2007) instead of SW winds (this study). These seemingly contradictory observations may not be in conflict if prevailing SW winds during the summer, like the modern winds, shaped the front of the Provo delta of the Bear River and the later parabolic dunes, but stronger winter storms from the NW were ineffective in northern Cache Valley due to shallow water with small fetch, flanking highlands, and the presence of abundant winter ice across the shallow lake surface in northern Cache Valley. Other Unresolved Topics Still to Understand 1. The Riverdale fault zone and whether it is a normal fault that ruptured beneath Lake Bonneville, whether it ruptured after the flood, and whether it is perhaps the headscarp of a large lateral spread. 2. The age of the Riverdale structure is critical to determining the trigger for the Bonneville flood. 3. The upper limit of scours of the Bonneville flood, at ~4921 ft (1500 m) in northern Cache Valley, is ~50 m below the Bonneville highstand shoreline. 4. The sill for the higher Provo shoreline at 4775 ft (1456 m) is not preserved, and was probably removed by erosion. This makes it more difficult to determine its original location.

5. Details of the reversal of drainage directions in northern Cache Valley since late Provo time. CONCLUSIONS Analysis of landforms on aerial photographs, digitalelevation models, and satellite imagery, consideration of gravity data, analysis of drillers’ logs of water wells, and geologic mapping show that complex relationships in northern Cache Valley and southern Marsh Valley differ significantly from the commonly accepted history of Lake Bonneville. A resistant conglomerate in the Salt Lake Formation and the broad dam-like morphology and northwest slope of the Marsh Creek pediment and alluvial fan may have allowed episodic overflow without significant downcutting prior to the failure of the Zenda threshold. The Bonneville flood may have been triggered by an earthquake on the Riverdale fault, a landslide that had its headscarp along the fault zone, or by seiche waves generated by faulting and/or landslide failure. Other processes probably also contributed to the timing and geometry of the flood. During the Bonneville flood, the main outflow channel first occupied and cut a curving channel on the eastern side of Red Rock Butte. Later, the flood was diverted entirely into a deeper, straighter channel on the western side of Red Rock Butte, perhaps by a landslide that redirected flood-related incision. Modest but repeated landslides collapsed from hillslopes directly west and east of the scour channel of the Bonneville flood near Red Rock Pass. Voluminous flood water re-excavated scour channels filled by landslides, removed large volumes of bedrock during scouring, and eventually created a fluted and scoured landscape beneath the northern 11 km of the former lake basin of Cache Valley and 14 km of southernmost Marsh Valley. Scours removed soils and loose materials from surfaces in the former lake basin below 4921 ft (1500 m). This altitude is ~177 ft (50 m) below the altitude of the lake when it failed. Scours are only present within the Swan Lake terrain, north of an east-northeast–trending bedrock ridge through Swan Lake, and are notably absent farther south. The scours formed only on and north of the threshold for the higher Provo shoreline of Lake Bonneville. Wholesale failure of the northern Bannock Range in a megalandslide did not occur. The southwest-dipping Riverdale normal fault disrupts downslope gullies in foresets of the Bonneville delta of the Bear River, and likely was coeval with the Bonneville flood. A moderate to large earthquake on this fault might have triggered the Bonneville flood. Topsets of the highstand Bonneville delta and the two Provolevel deltas of the Bear River formed subaqueously at lower altitudes than their associated shorelines. Subaerial parts of the Bonneville delta were upstream in Gentile Valley and within Oneida Narrows. The Bonneville delta of the Bear River was abandoned during the Bonneville flood ca. 17.4 ka. Its foresets and distal topsets were extensively dissected during and immediately after the Bonneville flood by numerous closely spaced, parallel, downslope-trending gullies that excavated 20–45 m of sediment.

Reinterpreted history of latest Pleistocene Lake Bonneville The more widespread Provo delta of the Bear River filled all of northern Cache Valley from east to west with fine sediment, and it is barely dissected by younger incised tributaries of the Bear River. The highstand Bonneville delta of the Bear River is slightly convex basinward, perhaps reflecting a stronger riverine influence, whereas the Provo delta is concave basinward, likely due to strong wave action. The erosional event that dissected the highstand Bonneville delta was short-lived, was probably unrelated to rapid dewatering after the Bonneville flood, and was perhaps due to stormier, wetter(?), windier, and more fluctuating erosive conditions of the late Pleistocene. Rapid exposure and initial erosion before colonization by plants also may provide an explanation for the marked contrast between the preservations of surfaces of the Bonneville and Provo deltas of the Bear River. There are two Provo shorelines in northern Cache Valley. These two lake levels were controlled by two separate bedrock sills. The northward outlet for the Bonneville River (Gilbert, 1880, 1890) that formed immediately after the Bonneville flood was along the Swan Lake scour channel, probably near Swan Lake, atop the NE-trending Swan Lake horst, in the footwall of the Riverdale fault. The Provo shoreline does not persist north of this location (Fig. 2). The Swan Lake sill controlled the level of Lake Bonneville at the more prominent older and higher 4775 ± 10 ft (1456 m) Provo shoreline. Gilbert (1890, 178) came close to predicting the existence of this bedrock sill when he observed that “…the outflowing (Bonneville) river headed…farther south, between Swan Lake and Round Valley Marsh.” The second, lower Provo shoreline, near 4745 ± 10 ft (1447 m), is more subtle than the higher Provo shoreline. This lower lake stand had a bedrock sill ~23 km south of the original Zenda threshold. We interpret this episode to be shorter than that for the higher shoreline, because a longer occupation would have produced a more definitive shoreline. The Bonneville River flowed north from a second, younger, and lower sill near Clifton, Idaho, into Round Valley. There it meandered laterally and produced point-bar scrolls and a large meander belt and floodplain while flowing northward toward the Snake River Plain. This area is occupied now by a marsh that flows sluggishly southward during high-water years. During regression of the lake below the lower Provo level, the entrenching Bear River built two successively lower, small deltas into northern Cache Valley. After the lake fell below the lower Provo level and the Clifton sill during the Holocene Marsh Creek built an alluvial dam into the Swan Lake scour channel at Red Rock Pass (Williams and Milligan, 1968). This established the modern drainage divide at Red Rock Pass, 2 km south of its original pre-flood position near Zenda, Idaho. This uneven infilling of the originally north-sloping dry bed of the Bonneville River reversed the flow direction between Clifton and Red Rock Pass. The coincidental repositioning of the modern drainage divide close to its original one near Zenda, Idaho, may explain why Gilbert (1880, 1890) was the only prior researcher to interpret a major southward shift of the outlet of Lake Bonneville during occupation of the Provo shoreline.

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DETAILED FIELD-TRIP LOG (First number = transit miles; second number = transit time in minutes; third number = total time in minutes.) Total Mi Min Min 0.0

0

0

1.2

3

3

4.3 11 2.7 5

14 19

2.9

5

24

3.0

4

28

2.7

4

32

0.9

1

33

1.7 1.2

3 4

36 40

Depart north on 100 East from east side of Riverwoods Convention Center. Utah State University sits atop the Provo delta, near 4780 ft altitude here, on the north side of the inset valley cut into the delta. The Logan Latter Day Saints Temple is on a small post-Provo delta remnant that formed during a brief still stand near 4640 ft altitude, whereas River Heights, east of the convention center, lies on a second small postProvo delta remnant near 4570 ft altitude. A small delta (?) remnant at the highstand Bonneville level, near 5135 ft, is present on the south side of Logan Canyon, but not on the north (it is the same as at Providence Canyon and Blacksmith Fork Canyon to the south in Cache Valley). This, and spits built mainly southward from headlands, are possible evidence of dominant southward longshore drift in Lake Bonneville. The East Cache fault, along the west base of the Bear River Range, the Provo delta foresets, and the river-cut slopes bounding the inset valley are three types of escarpments. Turn west on 400 North, then north on U.S. Highway 91. Stoplight at Hyde Park Lane. Stoplight, at 100 North at Center Street in Smithfield, is at the crest of a low post-Provo alluvial fan. Salt Lake Formation crops out on the west side of U.S. Highway 91. Stoplight is at the crest of a spur of Salt Lake Formation in Richmond, Utah. At the road to east to Cove, Utah. The valley of the Cub River, west of Highway 91, is inset ~40 ft below the lake-bottom surface of Lake Bonneville. Pass Utah Highway 61, which goes west to Lewiston. Idaho state line. Original La Tienda store in Franklin, Idaho, which figured in the movie, Napoleon Dynamite, about life at the Preston, Idaho, high school. The newer store still sells more Idaho lottery tickets than most other outlets in Idaho.

214 0.4 2.0

Janecke and Oaks 2 2

42 44

2.5

3

47

2.1

5

52

0.9

2

54

1.7

4

58

1.7

7

65

0.8

1

66

0.9

2

68

0.3

2

70

Cub River, tributary of the Bear River. Cub River Road, to east-northeast, joins U.S. Highway 91 where the latter bends to the northwest. Note the absence of a delta at the mouth of the large Cub River Canyon. Worm Creek is one of four deep post-Provo gullies through Provo delta foresets of the Bear River. This gully is inset below the Provo delta bottomsets southward, where it joins the Cub River. Here U.S. Highway 91 starts to rise up the foresets. U.S. Highway 91 turns north in Preston, well onto the topset of the Provo delta of the Bear River. U.S. Highway 91 turns west, where Idaho Highways 34 and 36 turn east-northeast, on Provo delta topset near the north edge of Preston Idaho. Preston Airport is at the crest of an erosional scarp ~50 ft high, formed by the Bear River during initial recession below the Provo topset. The lower surface to the north and west is likely a strath terrace corresponding to the post-Provo delta ~5 km south. There is a hot spring on the Bear River here north of the inflow of Deep Creek. Note the highest post-Provo shoreline level along the west flank of this entrenched valley. Numerous slumps, with common earthflow toes and arcuate headscarps (several historic), mark the steep valley walls where fine-grained sediments of the Provo and Bonneville deltas have failed due to undercutting by the meandering Bear River. The base of the river-cut scarp, with landslides, here lies at the north edge of the Bear River floodplain. Battle Creek, to the northwest, was site of the Bear River Massacre in 1863. That was the last major attack on Native Americans (Shoshone) by the U.S. Cavalry. Turn east on side road to Scenic Lookout, at the crest of the river-cut scarp. We have returned to the topset of the Provo delta of the Bear River. Stop 1: 4727 ft, 42° 09.176′ N and 111° 54.438′ W. This overview stop will be used to orient participants to the landscape. The large mountain ranges are obvious, with the Bear River Range in the east and the Bannock Range in the west. Our stop is at the extreme distal edge of the Bonneville delta of the Bear River and at the upslope edge of the Provo delta of the Bear River. The higher Provo

shoreline is a little bit east of us, but is difficult to identify in the field on the deltas. The Bonneville delta prograded halfway across this part of Cache Valley from the inlet of the Bear River south of Oneida Narrows, after filling that gorge and Gentile Valley just upstream. The landscape occupied by the Bonneville delta is subtle, with its deeply dissected foreset rising immediately east of us, the delta’s topset in the distance (~6 km away and ~100 m higher) and the Bonneville shoreline in the far distances etched across the Paleozoic rocks of low, distant foothills (7–8 km away). A view from overhead would provide a view of the several tens of gullies eroded into the Bonneville delta of the Bear River and the small number of widely spaced and much deeper gullies that cut into the Provo delta of the Bear River. Also we would see 2–3 deep gullies cut parallel to the Riverdale fault zone 3– 4 km east of us. This fault zone (or possible landslide?) does not disturb any Provo-age deposits along strike to the northwest but has a clear and obvious expression across the Bonneville delta (Fig. 11). The cross-cutting relationships show that the fault (landslide?) slipped right around the time of the Bonneville flood. More precise dating is needed to pin this down further. The flat high terrain underfoot, to the west, to the south of the Bear River, and to the southwest is the topset of the Provo delta of the Bear River. The incision by the modern Bear River is the main modification to this delta since its abandonment at the end of the last pluvial. We were on this topset in the Preston area and will continue driving across it to our next stop. The Provo delta of the Bear River is large, and filled the remainder of northern Cache Valley from east to west with fine sediment. The distal southern margin of the Provo delta is completely unlike the distal western margin of the Bonneville delta of the Bear River. The Provo delta is concave basinward, whereas the Bonneville delta is convex basinward. This geometry reflects a change from river-dominated deposition during the Bonneville highstand to wave-dominated deposition at the Provo still stand, at least at the delta front. We infer that this contrast in shape and the contrasting degrees of erosion are due to markedly different climates. The bedrock hills closest to us, in the west and northwest, are the Twin Lakes composite horst block with its faulted mix of Neogene Salt Lake Formation to Neoproterozoic bedrock (Link and LeFebre, 1983). This irregular set of hills pokes up through the topset of the Provo delta, and has two northern appendages that are buried just beneath the Provo sediment. The western buried bedrock ridge was the outlet of Lake Bonneville during the brief(?) stillstand at the lower 4745 ft Provo shoreline. The evidence for this interpretation will be discussed at Stop 8, in Round Valley on the far side of the Twin Lakes horst. We can see some of the sediment of the deltas in cuts along the Bear River Valley. No gravels are apparent in the steep scarps exposed by multiple landslides, and few are recorded in numerous drillers’ logs of water wells. The delta is composed of silt, clay, and some sand in the subsurface. The reddish color of this

Reinterpreted history of latest Pleistocene Lake Bonneville fine sediment is inherited from its Tertiary and Mesozoic source rocks in western Wyoming and SE Idaho. Inset remnants visible upstream along the Bear River at the Provo-delta level may be strath terraces in part. 0.2

2

92

4.0

6

98

6.4 10

108

Return to U.S. Highway 91. Turn north. Note the proximity of reddish sediment in the foreset of the Bonneville delta of the Bear River northeast of us. At pressurized water pipeline for irrigation, we cross a subtle saddle and buried bedrock of the Twin Lake horst that forms the southeast corner of Round Valley. Relict meanders of the north-flowing Bonneville River lie in the lowest part of the valley, west of us, as we drive north to Stop 2. Stop 2: 4796 ft, 42° 17.658′ N, 111° 59.225′ W. This stop is at a road cut on the east side of the highway and east of the south part of Swan Lake. Most of the exposure is covered because the bulk of the Salt Lake Formation is fine and tuffaceous. Locate the single poorly sorted conglomerate bed of the Salt Lake Formation in the upper half of the outcrop, and trace it laterally to see how it defines a broad, faulted anticline that trends roughly east-west. This anticline in the Salt Lake Formation projects west along the Swan Lake ridge and horst. The ridge coincides with widely scattered exposures of brecciated bedrock all the way up to the base of the Bannock Range in the distance (Long and Link, 2007), and overlies a pronounced gravity high (Kruger et al., 2003). Floodrelated scours sliced into the top of this ridge. Scoured hillslopes of the Bonneville flood end along the SSE edge of the Swan Lake bedrock ridge. The flood scraped deep groves into the Swan Lake horst block from south to north—just like a rake scrapes shallow grooves into moist sand.

Scoured landscapes are notably absent farther south (Figs. 7). Twin Lakes horst, just a few kilometers to the south of the scoured landscapes, has magnificent transgressive Bonneville shorelines stacked from top to bottom (some hillslopes preserve ~20 beach ridges) without evidence of later scour. This is striking because Twin Lakes horst is located in the center of the basin that was draining catastrophically during the Bonneville flood, and it lies in the same altitudinal range as the scoured landscapes in the Swan Lake terrain to the north. The contrast between the pervasively scoured landscape from the Swan Lake horst northward and the rest of the Bonneville basin farther south suggests that scouring developed pri-

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marily downstream of the new outlet of Lake Bonneville in the late stage of flooding. For scours of this intensity to form, the landscape must be near or within the outflow channel and thus must be north of the eventual sill for Lake Bonneville after the Bonneville flood. 1. The Swan Lake scour channel starts here and has the morphology of a megaflood-related scour (e.g., Baker, 2009). 2. Its geometry is most consistent with a sill at its south end, not one in the middle at Red Rock Pass (Fig. 10). 3. Notice the irregular topography SW of here. The flood scours might have produced this landscape, but its morphology also resembles a landslide mass and its poor, blocky exposures are similar to those northward toward Red Rock Pass. This likely was the outlet of Lake Bonneville at the older, higher Provo shoreline, when drainage of the Bonneville River persisted northward along Swan Lake scour channel. The Riverdale fault, downthrown to the southwest, lies buried ~3 km to the south. Swan Lake is a shallow lake that formed in the Holocene. It is a residual low between adjacent infilling alluvial fans, and drowns a part of the Swan Lake scour channel that was filled less than areas to the north and south (Fig. 7). A piston core (Bright, 1966) at Swan Lake showed 9.55 m (31.3 ft) of Holocene fill here, to ~4730 ft altitude, above Salt Lake Formation. Two water wells within this channel, 4.3 km NNW of this stop, showed two laterally correlatable alluvial gravels separated by 10–13 ft (3– 4 m) of clay, to a depth of 101 ft (31 m), near 4678 ft (1426 m) altitude. The base of the lower gravel is ~97 ft (30 m) and 67 ft (20 m) lower, respectively, than the 4775 and 4745 ft shorelines. Thus the sill probably was farther south. From this stop north we pass by a series of low hills and valleys. Irregular topography here has up to 500 ft (~150 m) of relief, and pre-Quaternary bedrock of all types is close to the surface in many places (Bright, 1963; Link, 1982a). 1.6

6

134

3.3

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Swan Lake hamlet and country store, is within the Swan Lake scour channel. Fluted terrain to the west, basinward and lower than east-sloping pediment and alluvial fan remnants, is visible to the northwest. Landslide blocks related to the Bonneville flood are common along our route north of here. Stop 3: 4888 ft, 42° 21.172′ N, 112° 02.656′ W. Turn right on Red Rock Road. Turn around at 0.1 mi. Park at base of bedrock hill west of the road. Climb to hillcrest for a view of the scoured landscape produced by the Bonneville flood. North of us is Red Rock Butte, and west of us is Red Rock Pass (~under the overpass). Red Rock Butte to the north consists of Cambrian to Ordovician St Charles Formation (Fig. 13; DeVecchio et al., 2003). Seismic refraction of Williams and Milligan (1968) suggested bedrock at a depth

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Janecke and Oaks of ~6 m (20 ft), near altitude 4755 ft immediately north of the overpass. However, the distinct refraction in their study might have imaged the top of landslide debris.

North of Stop 3 are two scour channels that formed sequentially (?) during the Bonneville flood. The eastern channel, east of Red Rock Butte, has more curvature, formed during the Bonneville highstand, and was active during the early part of the flood. The east scour cut across older Pleistocene sediments along the east wall (Stop 4), but the preserved part of its west wall is durable Paleozoic rocks, and the floor exposes tuffaceous Salt Lake Formation. The Pleistocene sediments vary in thickness. The field relationships may show that downcutting slowed or stopped as more resistant rocks were being exposed in the base of the scour (at an altitude of ~4870 ft [1485 ± 1 m; Fig. 5]). However, some other process probably deactivated this channel, because the tuffaceous facies of the Salt Lake Formation here is weak. The landslide beneath our feet, or an older one in a similar location may have been responsible for deactivating the eastern scour and for diverting flow entirely into the western scour channel (Fig. 13). The straighter, lower, western scour channel became the only locus for outflow of Lake Bonneville during the last ~50 m of incision in this area. This western channel is better known than the eastern one, and might have been active from the start of the Bonneville flood. Our reconstruction in Figure 13, however, shows another alternative, that the eastern scour was the sole locus of outflow early during the flood and that the western scour replaced it later in the flood. The western scour became the bed of the Bonneville River, the very large but short-lived north-flowing river that emerged from Lake Bonneville when it was overtopped (Gilbert, 1890). The lake was hydrologically open northward when the Provo shorelines were occupied between ca. 17.4 and 15 cal ka BCE (Godsey et al., 2005). We are standing on a block of Cambrian Blacksmith Formation within a large landslide that was shed westward (DeVecchio et al., 2003) into the scoured area. Notice that these Paleozoic rocks dip ~30° toward the east-southeast, toward the headscarp in the east. This is one of several prominent landslides of the Red Rock Pass area. We can see other landslides as well from this stop and there are erosional remnants in the scour channel and above the west edge of the western scour channel. We mapped these landslides, and found that they are localized within 1–2 km of the scour channel and are most abundant in the foothills of the Bannock Range (Fig. 13). The area affected by landsliding in the northern Bannock Range (Fig. 4) is much smaller than the >17 km2 megalandslide hypothesized by Sewell (cited in Smith et al., 1989). Large areas that Sewell interpreted as landslide are instead stable pediment remnants, distal fan deposits and exposures of in-place Salt Lake Formation (Mayer, 1979; Long and Link, 2007). The west channel bottom is nearly flat due to post-Provo infilling with sediments by the Holocene Marsh Creek (Gilbert,

1890; Bright and Ore, 1987), whereas the east channel was abandoned during the flood, and is entrenched by Marsh Creek. The east channel was not significantly infilled. The modern drainage divide is localized by the slight rise on the radially bifurcated Holocene Marsh Creek alluvial fan (Figs. 4, 8, and 10). 0.1

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Return to U.S. Highway 91, turn northwest, then turn northeast onto dirt road just before overpass bridge. Marsh Creek. We are driving on the northeast side of bedrock promontory between higher east Bonneville flood channel and lower west Bonneville flood channel. Fan-pediment remnant to the east here is truncated by, and forms the east wall of, the higher flood channel. The surface northward here is the abandoned floor of the east flood channel. Where the road begins to descend, the northwest edge of the bottom of the east flood channel was eroded to form the east edge of the west flood channel. Pratt Road. Turn northeast and drive across the eastern scour channel. Continue and stop at the outcrops on the right after starting to drive up a gully that is cut into the east wall of the channel. Stop 4: 4935 ft, 42° 22.535′ N, 112° 02.880′ W. The east scour channel, to our west, is cut into the pale, gently inclined sand and gravel on the south (right) side of the road. These Quaternary sediments are probably the distal deposit of the Pleistocene Marsh Creek alluvial fan (DeVecchio et al., 2003; Fig. 6). Characterized by well-bedded, sorted sand, silt, and gravel with rounded clasts, this is probably a fluvial deposit. These deposits are unconsolidated but cohesive, and contain thin gravel beds and sand. They likely had weak resistance to floodwaters. The eastern scour channel cut through this sediment, and stopped downcutting when it encountered the more indurated Salt Lake Formation below.

This pre-Bonneville, Marsh Creek alluvium contrasts with the darker, more orange sediment in the west part of this outcrop. These orange sediments are also Quaternary and overlie the fan deposits along a buttress unconformity that dips west toward the channel. One large boulder of the fan sediment was incorporated into the darker sediment and documents the relative age as younger than the alluvium and younger than the eastern scour. The darker sediment is poorly sorted and weakly bedded sand and gravel. It might be an unusually thick hillslope deposit that was shed from the oversteepened east wall of the eastern scour channel. The grain size seems too fine for these to be flood deposit.

Reinterpreted history of latest Pleistocene Lake Bonneville This Marsh Creek alluvium and its overlying inclined sediment are in a cut-and-fill relationship with underlying Salt Lake Formation (Stop 6), and thin northward. Be sure to examine the thin gravel lenses for their clasts. We will compare the clasts in these Quaternary gravels with those in Miocene conglomerates at Stop 6. Continue driving east. 0.8 1.0 1.7

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Head of gully, road bends to north. At road junction, turn west. Stop 5: 5131 ft, 42° 24.085′ N, 112° 04.273′ W. Proceed west on the gravel road until you are at the top of a west-facing hillslope. We can see the large, open drainages produced by sapping below the highest altitude of groundwater seeping north from Lake Bonneville. Scalloped drainages with theaters at their heads and margins persist all the way to the valley floor to the northwest, but end upslope a few meters above the altitude of the Bonneville highstand shoreline. Higher on this pediment, dry gullies have V-shaped cross sections typical of fluvial erosion. This groundwater flow and sapping could explain the Bonneville flood, if the basin was closed and sapping weakened the sill and caused it to fail (O’Connor, 1993).

Our vantage at Stop 5 allows us to see the pediments and alluvial fans along the southern flanks of Marsh Valley. The remnants of two large fan-pediments here truncate folded and faulted Neogene Salt Lake Formation, and head in Aspen Creek (west) and in Marsh Creek (east) (Fig. 4). These were graded to a common base level that is well above the modern streams. Northern Cache Valley contains some smaller remnants of pediments and alluvial fans that graded to a similar higher base level. The Bonneville shoreline is cut lightly into part of this pediment east of Red Rock Pass (Gilbert, 1890) and in northern Cache Valley, so the pediments and Pleistocene alluvial fans exemplify the landscape before the flood. We used the regular geometry of these surfaces to reconstruct the landscape to a pre-flood condition, and found that the reconstructed saddle was coincident with the highest shoreline of Lake Bonneville, within errors. Therefore, we agree with most prior workers that Lake Bonneville probably had an outlet and was an open lake when it reached its highest altitude at the Bonneville shoreline (cf. Currey, 1982; Currey and Burr, 1988; Currey, 1990; Oviatt, 1997; Oviatt et al., 1992; Currey and Oviatt, 1985). A sapping-related dam failure roughly at the altitude of the highest Bonneville shoreline is an ineffective way to start a catastrophic flood if, as we believe, Lake Bonneville was intermittently flowing out across the Zenda pediment threshold prior to the flood (see also Currey, 1982, and others). The broad, gentle curvature of the surface of the coalesced pediments, their gradual downstream slope, and a resistant conglomerate within the Salt

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Lake Formation near the low point between the pediments (Stop 6) apparently made this area a surprisingly long-lived (~1.2 k.y.; Godsey et al., 2005) earthen dam for Lake Bonneville. Note the pediment-covering sediment on Salt Lake Formation across the valley, to the southwest (toe near 5060 ft, rising to ~5250 ft). In distal settings and in cut-and-fill locations, the Quaternary sediment is thicker than the pediment-covering deposits. 3.9 10

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Turn around, return to junction, turn south. Return to Pratt Road junction, and turn north. Note ash in Salt Lake Formation in road cut here. This is late Miocene ash from the Twin Falls caldera along the Yellowstone hotspot track (DeVecchio, 2002; DeVecchio et al., 2003). At Zenda, Idaho (a few ranch houses) turn west. Note pediment on Salt Lake Formation to west-northwest, with landslides along undercut face in front. U.S. Highway 91. Turn north. Entrance to Downata Hot Springs resort. Turn west. Downata Hot Springs main building. Park, use facilities, and gather at picnic tables east of the main building for lunch. Return to U.S. Highway 91, turn south, proceed 0.9 mi.

Stop 6: 4764 ft, 42° 23.021′ N, 112° 04.172′ W. Park on west side of road at base of a small hill. Inspect the blocks of cemented conglomerate in the road cut and then climb to the top of the hill for a view of the scoured landscape very close to the original threshold of Lake Bonneville before the flood. The Zenda dam would have been ~90 m above us and ~800 m south of here before the Bonneville flood. The flood initiated here, in these fairly weak sedimentary rocks with thin overlying pediment-related Quaternary sediment and loess, after holding fast for long enough to produce a well-defined shoreline across the Bonneville basin. We estimate a minimum duration of at least 500 years for the Bonneville highstand because the ~2000-year occupation of the higher Provo shoreline (Godsey et al., 2005) produced a shoreline of similar intensity. At Stop 7 we will show that an earthquake might have triggered the Bonneville flood. The hill under our feet is an erosional remnant of a landslide that was shed west into the valley from the oversteepened east wall of the Swan Lake scour and discharge. The landslide was mostly eroded away by floodwaters, by the subsequent Bonneville River, or both. The blocks of cemented conglomerate in this landslide were carried downslope from the in-place conglomeratic Salt Lake Formation from the hillside directly east of us, where it is less accessible. Beds of silt and fine sand alternate with conglomerate and gravel along the entire eastern margin of the scoured channel here, and formed the dam for Lake Bonneville at its highest level. The degree of cementation varies somewhat, but

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these blocks are typical of conglomerates exposed in the channel wall from Zenda northward to the end of the exposures. Several features show that this landslide transported conglomerates of the Neogene Salt Lake Formation and not Quaternary alluvial-fan deposits, as originally mapped (DeVecchio et al., 2003). In particular, the clast composition is very different from that in the distal alluvial sediment of Marsh Creek exposed at Stop 4, in having many conspicuous reddish and purplish quartzite clasts derived from the Neoproterozoic Brigham Group in the Salt Lake Formation (as we can see at this stop) in contrast to a few white quartzite and much dark to black chert and carbonate clasts in the Marsh Creek alluvial-fan deposits (at Stop 4). Dips in the Salt Lake Formation are low but toward the east throughout this outcrop belt (DeVecchio et al., 2003; Janecke, unpublished mapping). Finally, this conglomerate of the Salt Lake Formation resembles conglomerate beds in a strongly tilted and later eroded subunit of Salt Lake Formation that is interbedded with thick white and silver tuffs in the west wall of the Swan Lake scour channel near Downata Hot Springs (our lunch stop). Ashes that are downsection of the in-place conglomerate, south of our stop, were chemically correlated to 7–10 Ma ashes erupted from the Yellowstone hot spot track (DeVecchio, 2002; DeVecchio et al., 2003). 2.4

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Return south on U.S. Highway 91 to approach to overpass bridge that crosses the railroad tracks. Turn south on Idaho road “D 1.” Note the flat-bottomed and infilled Swan Lake channel to the south-southeast with numerous knobs of flood-sculpted landslide material sticking up. The reddish weathering cavernous cliff of Cambrian limestone west of the road was transported in one of the many modest-scale landslides that initiated during the Bonneville flood (Fig. 13). We speculate that reservoir-induced seismicity from the rise and fall of Lake Bonneville may have helped to destabilize some of the hillslopes during and after the flood.

We will drive south parallel to high remnants of pediments and alluvial fans that are truncated at their toes by multiple east-facing escarpments that could be flood scours, fault scarps, or both. We suspect that most of these are fault scarps of the Dayton-Oxford fault zone because they project into other probable faults, and flood-related scours are lower in elevation, below ~1500 m. The Bonneville flood stripped the landscape of its thick darker soil within all the other scours, thereby exposing light-colored tuffaceous Salt Lake Formation or light preTertiary bedrock. 4.3

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Stop 7: 4965 ft, 42° 16.507′ N, 112° 00.720′ W. The stop is located near the southern margin of the Swan Lake horst block, overlooking Round Valley, south of us. Our location is within an area that was scoured and modified by the Bonneville flood to form N-S trending scours and ridges. A short distance west of here there is no clear evidence for scouring in the 50 m below the high Bonneville shoreline. Meanders are visible in the lowest part of Round Valley below us to the south (Fig. 9). These large relict meanders have wavelengths of 2.3 km, and match the size of those produced by the Pleistocene Bear River (Fig. 12). The meander belt is at an altitude of 4745 ft (1446 m), and is distinctly below the higher main prominent Provo shoreline in northern Cache Valley, the 4775 ft (1456 m) shoreline. Therefore a separate, lower, and probably later Provo shoreline is indicated by these relationships. We will examine the lower 4745 ft (1446 m) shoreline at Stop 9 in the Weston, Idaho, area. 0.3

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Return to Idaho road D-1, turn left; note road crest ~4948 ft at road cut in Salt Lake Formation just southeast of well at Stop 7. Quartzrich rock was encountered at 82 ft in the well at the home on the hill east of the road. Road bends to southwest at canal. Note view southeast of formerly north-flowing meanders in Round Valley. Between here and Clifton, cobbles and boulders of foliated green rocks are common at the surface. They are mostly phyllite, derived from the Neoproterozoic Pocatello Formation in the cliff faces to the west. In a well log at 7375 North Westside Highway, a surficial unit composed of this material 29 ft thick, perhaps a landslide deposit, was identified as “basalt.” Clifton, Idaho. Turn right (west) and drive a short distance across a canal and stop at entrance to a small sand and gravel pit for an overview of the Clifton sill area.

Stop 8: 4766 ft, 42° 10.017′ N, 111° 59.828′ W. Stop and look around at the locked gate of a small borrow pit, west of the west side canal. In the gravel pit and to our west is faulted Salt Lake Formation (the white and silver sediment) in fault contact with Neoproterozoic metasedimentary rocks (Carney and Janecke, 2005; Carney et al., 2003). These rocks are tilt bocks in the hanging wall of the Bannock detachment fault. The Twin Lakes composite horst is the low set of hills east of here, and they separate Round Valley from the rest of Cache Valley. These low hills are largely between the Provo and Bonneville shoreline, and have closely spaced beach ridges along many of

Reinterpreted history of latest Pleistocene Lake Bonneville their exposed hillslopes. There is a thin carapace of these lake beds on complexly faulted Neoproterozoic Pocatello Formation, Brigham Group quartzite, Paleozoic carbonates, and much Late Cenozoic Salt Lake Formation (Link and LeFebre, 1983). This Twin Lakes composite horst block consists of two crossing fault blocks in a T-shaped geometry overall, with a chubby P-shape above ground level. One sub-horst is a narrow NW-trending bedrock ridge of Proterozoic rocks (Link, 1982a, 1982b; Link and LeFebre, 1983). We infer the Clifton sill at the narrow valley here, where the Bannock Range and Twin Lake horst block come closest to one another. This unremarkable looking place separates northwardmerging tributaries in the north from southward-merging tributaries in the south. Gravity data show shallow bedrock beneath this area (Eversaul, 2004). Locations of sparse logs of water wells are poorly documented across this area. However, shale of the Salt Lake Formation was reached at 129 ft (39 m) near the west margin, and hard clay was reached at 136 ft (41 m) near the east margin. The top of the Salt Lake Formation may be higher in both wells. Incision by younger streams is dissecting the topset of the Provo delta of the Bear River in the south, but has not begun to incise north of the sill. This is one of several geomorphic contrasts across the sill. 4.3

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In south part of Dayton, at junction with Idaho Highway 36 West, turn southwest. Idaho Highway 36 turns west. Continue south. Note older, higher Provo topset on Weston delta, near 4782 ft altitude here. Stop 9: 4747 ft, 42° 2.245’N 111° 58.898’W. Curve to east, and descend to younger, lower Provo shoreline. Park on north side of Depot St., southwest of the Latter Day Saints church. The Dayton-Oxford fault lies west of Idaho Highway 36 northwest of Weston and forms the east face of Rattlesnake Ridge. The step down to the east located west of the church is the 4745 ft Provo shoreline, and it is cut into the topset of the Provo delta of Weston Creek. This stop and alternate Stop 9 show the 4745 ft shoreline in particularly well-expressed locations.

Alternate Stop 9: 42°01.449′ N 111° 58.807′ W. A better location for viewing both the higher and lower Provo shorelines is south of Weston on a gravel road at these coordinates; however this stop might not be used due to time constraints. Stop here on the topset to the Weston delta and look around to see the 4745 ft shoreline due west of you at 42° 01.461′ N 111° 58.897′ W. This is below the main higher 4775 ft shoreline visible on the hillslopes to the south. The two shorelines are 52–56 ft (16– 17 m) apart vertically in this location. The Bonneville shoreline is also clear southwest of this location along the northeast flank of Bergeson Hill. Notice that the hillslope between the Bonne-

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ville and higher Provo shoreline is riddled with small steep gullies that trend downslope, gullies are closely spaced, and this intense erosion end downslope at the higher Provo shoreline with its distinctive wave-cut cliff and basinward bench. These gullies cut into Salt Lake Formation and its fairly thin Quaternary cover. 3.5

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Continue east, then south to Utah state line. Road becomes Utah Highway 23. To west are Big and Bergeson Hills, which contain faulted Salt Lake Formation, an anticline near the crest of Bergeson Hill, and local eastdipping Ordovician Garden City Formation and brecciated Paleozoic rocks near normal faults. Farther south, note shorelines in the west, and, past Trenton, black OrdovicianSilurian Fish Haven Formation dipping ~75° east, but dipping west with minor underlying Eureka Quartzite near the Provo shoreline north of the borrow pit that is ~0.5 mi south of 9400 North. Then gray quartzite (Neoproterozoic?) dipping west appears, near the Provo shoreline, south of the borrow pit and a distinct NW-striking gully. All pre-Cenozoic rocks are overlain by Miocene-Pliocene Salt Lake Formation in a north-plunging syncline along the crest of Newton Hill (which is also called Little Mountain). Stop 10: 4495 ft, 41°52.501′ N, 111° 56.889′ W. Pull into old county road–gravel storage area east of Utah Highway 23, at bend. East-west geologic section with logs of 3 water wells and 2 oil wells encountered thin Quaternary deposits over thick Salt Lake Formation over possible Eocene Wasatch Formation over white quartzite over Neoproterozoic metasedimentary rocks, the latter near 6200 ft (east) and 5200 ft (west), perhaps evidence for extent of the Bannock detachment fault south to this latitude. Second (south) gravel pit has abundant and varied Neoproterozoic quartzite with no obvious source, distributed in two foreset sequences separated by a welldeveloped buried paleosol. Relations to the west with a gently south-sloping gravel bench (~4830 ft in south rising to ~4980 ft in north), between the highest Bonneville shoreline and the Provo shorelines, are unclear. However, the gravels below the paleosol may belong to the older Pokes Point or Little Valley lake cycle, and those above probably belong to the Bonneville or perhaps to the Little Valley lake cycle if it reached this height (Scott et al., 1983; McCoy, 1987; Oviatt et al., 1987;

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Oviatt et al., 1999). Dating by optical stimulated luminescence is in progress. Continue south on Utah Highway 23 to near crest. Turn south on 4800 West. Take care crossing the northeast-trending railroad track to south. Note Cutler Narrows to west, where Bear River exits Cache Valley westward. There were four sites along the Cache Butte Divide where water could flow from west to east at the Bonneville highstand, at the start of the Bonneville flood, but only one, Cutler Narrows, when the lake level fell below ~5000 ft altitude (Maw, 1968; Oviatt, 1986a, 1986b). Utah Highway 218. Turn east toward Smithfield. Cross buried trace of the DaytonOxford fault close to here. Flat valley floor consists of bottom set muddy deposits of Lake Bonneville. Road begins to rise from ~4420 ft to ~4450 ft, near east end of the sewage lagoons, onto a sand levee built by Bear River atop the lakebottom sediments. We cross the Bear River east of the Amalga cheese plant (north). An oil well drilled there reached carbonate bedrock between 5200 and 5500 ft (total depth). Small solifluction lobes form each spring in the sandy levee slopes along the Bear River here. Smithfield town center. Turn south on U.S. Highway 91 at stoplight. Arrive on the west side of Riverwoods Convention Center ~5–6 p.m.

ACKNOWLEDGMENTS Reviews of manuscripts related to this topic by Jack Oviatt, Jim O’Connor, Jim Evans, and an anonymous reviewer helped to clarify our interpretation of Lake Bonneville. The “Can humans cause earthquakes?” podcast of “Stuff you should know” led to our speculative correlation of the Bonneville flood and induced seismicity. Several landowners helped in locating water wells. We thank Jim Evans, Thad Erickson, and Lu Oaks for their help and support. REFERENCES CITED Anderson, S.A., 1998, Sedimentology, hydrogeology, and sequence stratigraphy of Pleistocene Bear River delta, Cache Valley, Idaho [M.S. thesis]: Pocatello, Idaho, USA, Idaho State University, 61 p. Anderson, S.A., and Link, P.K., 1998, Lake Bonneville sequence stratigraphy, Pleistocene Bear River Delta, Cache Valley, Idaho, in Pitman, J.K., and Carroll, A.R., eds., Modern and Ancient Lake Systems: Utah Geological Association Guidebook 26, p. 91–104. Baker, V.R., 2009, The Channeled Scabland: A Retrospective: Annual Review of Earth and Planetary Sciences, v. 37, p. 393–411, doi:10.1146/annurev .earth.061008.134726.

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