ABSTRACT

ORIGIN AND STRUCTURE OF THE POVERTY HILLS, OWENS VALLEY FAULT ZONE, OWENS VALLEY, CALIFORNIA

by Tatia R. Taylor

Structural field mapping and analyses of the Poverty Hills reveal a complex deformational geometry consistent with that of a transpressional positive-flower-structure uplift. Kinematic indicators show vertical and horizontal translation facilitated by reverse faulting with a strong oblique strike-slip component. The onset of deformation in the Poverty Hills is estimated at ~1.7 Ma, and the Owens Valley fault is proposed to have stepped westwards in kinematic response to the rapid development of regional transtensional fault patterns. Geometric modeling of uplift points to a constant slip rate of 1.27-1.44 mm/yr since ~1.7 Ma on the northern Owens Valley fault. These kinematic, geometric, and temporal relations are utilized to better constrain slip partitioning and distribution across the major fault systems in the region and to characterize the spatial and temporal evolution of the Owens Valley basin, the Owens Valley fault, and the Eastern California Shear Zone in this region.

ORIGIN AND STRUCTURE OF THE POVERTY HILLS, OWENS VALLEY FAULT ZONE, OWENS VALLEY, CALIFORNIA

A Thesis

Submitted to the

Faculty of Miami University

in partial fulfillment of

the requirements for the degree of

Master of Science

Department of Geology

by

Tatia R. Taylor

Miami University

Oxford, OH

2002

Advisor: Dr. Yildirim Dilek______

Reader: Dr. Brian Currie______Table of Contents

Page

TITLE PAGE...... i

TABLE OF CONTENTS...... ii

LIST OF FIGURES ...... iii

LIST OF TABLES...... vii

ACKNOWLEDGMENTS ...... viii

CHAPTER 1: GEOLOGIC BACKGROUND...... 1

CHAPTER 2: FIELD OBSERVATIONS ...... 10

CHAPTER 3: STRUCTURAL ANALYSIS ...... 25

CHAPTER 4: IMPLICATIONS ...... 28

REFERENCES ...... 85

ii LIST OF FIGURES

Fig. 1. Schematic map of geologic features in the Owens Valley region. 41 Modified from Martel et al., 1987.

Fig. 2. Lithologic map of Owens Valley region. From Beard and 42 Glazner, 1995.

Fig. 3. The Inyo Thrust and related structures. Modified from Stevens 43 and Olson, 1972.

Fig. 4. Map of late Cenozoic faults of the Eastern California Shear 44 Zone from the Garlock Fault to the Walker Lane.

Fig. 5. Relative motion of the Sierra Nevada with respect to the Colorado 45 Plateau since 36 Ma. Northwesterly motion of SN started ~10 Ma. From Wernicke and Snow, 1998.

Fig. 6. Faults of the northern Eastern California Shear Zone. 46 Adapted from Lee et al., 2001.

Fig. 7. Map of Owens Valley with bounding ranges. Location of 47 northern and southern grabens shown on east side of basin. From Gillespie, 1991.

Fig. 8. Enhanced Landsat 7 image of Poverty Hills and adjacent 48 geologic and geographic features. Ages given for dated volcanics of Big Pinevolcanic field.

Fig. 9. View to the east of the Poverty Hills. 49

Fig. 10. Structural and geologic map of the Poverty Hills compiled from 50 structural mapping and field-collected data. Taylor, 2002.

Fig. 11. Slickenlines on granodiorite boulder (not in place). East central 51 Poverty Hills.

Fig. 12. Fluvial and colluvial deposits on NW margin of Poverty 52 Hills, ~8-10 m above Tinemaha Creek.

Fig. 13. Angular unconformity in Plio-Pleistocene lacustrine sediments. 53 Eastern fault block near Tinemaha Reservoir.

Fig. 14. Map of Big Pine Volcanic Field with eruption ages of dated 54

iii volcanics and local faults.

Fig. 15. Shaft at uplifted contact between granodiorite and Paleozoic 55 rocks. Fault breccia of W-bounding normal fault on right. View to E.

Fig. 16. Structural and geologic cross section A-A′ (Fig. 10), 56 Poverty Hills.

Fig. 17. Schematic map of Owens Valley fault zone showing 57 orientations of regional and local principal stresses.

Fig. 18. Shear zone in granodiorite. East side Poverty Hills. 58 View to the north.

Fig. 19. Models of transpressional push-ups, or positive flower 59 structures. From Sylvester, 1988.

Fig. 20. Breccia at fault contact. Mississippian shale clasts in 60 CaCO3 matrix. Southeastern Poverty Hills.

Fig. 21. Tilted Quaternary basalt flow capping deformed 61 granodiorite on the north edge of Poverty Hills. View to the southeast.

Fig. 22. Oblique reverse fault in Mississippian shale. SW 62 margin of Poverty Hills. View to the northeast.

Fig. 23. Mississippian shale thrust above Pennsylvanian siltstone. 63 SE side Poverty Hills. View to NW. Jake for scale.

Fig. 24. Stacked wedges of sheared and faulted granodiorite. SE 64 Poverty Hills.

Fig. 25. Elliptical sag pond playa in structural half-graben. 65 Central Poverty Hills. View to the W.

Fig. 26. NW and W margin of Poverty Hills showing scarp 66 of W-bounding normal fault. View to SE.

Fig. 27. NE- and SW-dipping ‘limbs’ of columnar basalt sheet 67 stepping down to the SW across NW-trending basement steps. View to the SE.

Fig. 28. View north from northwestern Poverty Hills. 68

iv Fig. 29. Contoured equal area stereographic projection of faults 69 in granodiorite.

Fig. 30. Equal area stereographic projection of faults with slickenlines 69 in granodiorite.

Fig. 31. Contoured equal area stereographic projection of faults 70 in metasediments.

Fig. 32. Equal area stereographic projection of faults with slickenlines 70 in metasediments.

Fig. 33. Contoured equal area stereographic projection of foliations 71 in granodiorite.

Fig. 34. Equal area stereographic projection of foliations with 71 lineations in granodiorite.

Fig. 35. Contoured equal area stereographic projection of foliations 72 in metasediments.

Fig. 36. Equal area stereographic projection of foliations with 72 lineations in metasediments.

Fig. 37. Equal area stereographic projection of bedding in metasediments 73 in west central Poverty Hills.

Fig. 38. Distribution of fold axes in metasediments of the west 73 central Poverty Hills.

Fig. 39. Equal area stereographic projection of bedding in metasediments 74 of isolated hill in southwestern Poverty Hills.

Fig. 40. Distribution of fold axes in metasediments of isolated hill in 74 southwestern Poverty Hills.

Fig. 41. Equal area stereographic projection of bedding in metasediments 75 of the southern Poverty Hills.

Fig. 42. Distribution of fold axes in metasediments of the southern 75 Poverty Hills.

Fig. 43. Distribution of fold axes in metasediments of all domains. 76

Fig. 44. Equal area stereographic projection of bedding of sediments 76 and basalt flow east of Hwy. 395.

v

Fig. 45. Tilt restoration of younger (pink) and older (blue) 77 fluvial/lacustrine beds of tilted fault block.

Fig. 46. Equal area stereographic projection of ‘limbs’ (Fig. 27) of 77 disturbed basalt flow in southern Poverty Hills.

Fig. 47. Tilt restoration of basalt flow (blue) with directional flow 78 lineations (magenta) in the southern Poverty Hills.

Fig. 48. Spatial and temporal model of NW-directed oblique transtension. 79 Owens Valley basin prior to ~1.7 Ma.

Fig. 49. Spatial and temporal model of NW-directed oblique transtension. 80 Present Owens Valley basin.

Fig. 50. Slip distribution with calculated (this study) and published slip 81 rates for faults in the vicinity of the Poverty Hills and Owens Valley.

vi LIST OF TABLES

Table 1: Ages of dated BPVF basalts 82

Table 2: Summary Geochronology 83

Table 3: Offset Estimates 84

vii Acknowledgments I wish to acknowledge Dr. Yildirim Dilek for presenting me with the opportunity to pursue my project. My infinite gratitude goes to my son Jacob (who does not remember me as anything but his Mom, the student) for following me all over the country and providing the daily reminder that we have to pick our battles carefully and above all, life is far too short to get too serious about. I could not have finished without the help of some very important people. The faculty and graduate students at Miami University have given support and encouragement and sometimes just an ear when I really needed it. Shawn Irvin, Carrie Wright, and Art Losey will always have a couch in my living room. Darin Snyder has made time for me even when he had none (especially if it meant using Corel). Dr. Brian Currie was my reader and got my basemaps ready. Cathy Edwards went far beyond the call of duty so that I was able to travel and participate, and still know that things were safe at home. Thank you all. My friends and colleagues at the University of California White Mountain Research Station have served as a wealth of information and encouragement, and have made their home my home. Dr. Nancye Dawers, Tulane University, Dr. Jim McClain, University of California Davis, Dr. Anjela Jayco, U.S. Geological Survey, Menlo Park, and Dr. Clemens Nelson (ret.), University of California Los Angeles, and many other students and researchers, have all been an inspiration. Cheryl Seath at the California Bureau of Land Management in Bishop, the staff at WMRS, my field assistants Charlie Angerman, Jake Taylor and Chuck Taylor, and of course all the dogs at the station, made fieldwork in the blistering California desert a pleasure. Funding for this project came in the form of grants I received from the Geological Society of America and the University of California White Mountain Research Station, and additional support came from the Miami University Department of Geology and Dr. Dilek. This project was very important to me both personally and scientifically, and I am grateful for the means by which to pursue it and hopefully contribute some small piece of the puzzle.

viii CHAPTER 1: Geologic Background

Owens Valley The Owens Valley in southeastern California is the westernmost expression of Basin and Range extension between latitudes 36° and 37°30′ N, bounded on the west by the Sierra Nevada Mountains and on the east by the White-Inyo range (Fig. 1). The ~130 km long valley trends NNW, ranges in width from 10-22 km, and exhibits an elevation difference of ~3600 m with respect to the bounding ranges.

Sierra Nevada and White-Inyo Mountains The ranges flanking the Owens Valley are comprised predominantly of pre- Cenozoic rocks, while within the valley itself few exposures of basement rock exist among the extensive deposits of Cenozoic basin fill (Fig. 2; Pakiser, 1964). Rocks of the White-Inyo Mountains are highly folded and faulted (Knopf, 1918) Precambrian to Permian sedimentary and metasedimentary units which have been subsequently intruded by Mesozoic granitic plutons associated with the Sierra Nevada batholith (Knopf, 1918; Pakiser, 1964; Bateman, 1965). The Sierra Nevada Mountains to the west consist of a composite of Mesozoic granitic batholithic rocks which include roof-pendants of metamorphosed Paleozoic and Mesozoic units (Pakiser, 1964; Bateman, 1965; Stevens et al., 2000). The latest episode of pluton emplacement is punctuated by the 148 Ma (Late Jurassic) Independence dike swarm which divides older Sierra Nevada rocks from younger, post-148 Ma (mainly Cretaceous) Sierra Nevada plutons (Chen and Moore, 1979; Glazner et al., 1999; Coleman et al., 2000; Moore and Hopson, 1961). Within Owens Valley, isolated exposures of the pre-Cenozoic metasedimentary and plutonic basement derived from both the Sierra Nevada and the White-Inyo Mountains occur primarily in the Alabama Hills east of Lone Pine and the Poverty Hills.

Tectonic Geomorphology in Owens Valley Tectonic and geomorphic processes have affected the structure and stratigraphy of the volcanic, sedimentary, fluvial/lacustrine, and glacial rocks within Owens Valley.

1 Basin fill of the Owens Valley is a complex combination of alluvial sediments, alluvial fans and fanglomerates, paleo- and modern lake deposits, and volcanics. All have been faulted extensively, recording the tectonic evolution of the basin from the uplift of the bounding ranges in the mid to late Miocene through recent transtensional activity (Leuddecke, et al., 1998; Pinter and Keller, 1992, 1995; Berry, 1997).

Volcanism Volcanic rocks occur in three distinct areas: (1) the Coso volcanic field in the southern Owens Valley has been active periodically since the middle Miocene (Manley, et al., 2000) and most recently ~6 Ma (Duffield, et al., 1980); (2) the Volcanic Tableland north of Bishop resulted from the massive eruption and collapse of the Long Valley Caldera ~760 ka (Reheis and Dixon, 1996); and (3) the Quaternary Big Pine volcanic field (Fig. 1, 2) which surrounds the Poverty Hills dates from 1.2 Ma to as recent as 25 ka (Manley, et al., 2000; Martel, 1987; Duffield et al., 1980; Moore and Dodge, 1980; Bacon et al., 1982). All of the volcanics in the Big Pine field are potassic olivine basalts (Manley, et al., 2000). One anomaly within the Big Pine field is the presence of a small rhyolite dome just west of Poverty Hills that was dated at 1 Ma (Moore and Dodge, 1980; Mayo, 1944).

Active faulting and Sedimentation Evidence of active faulting is exceptionally preserved by the volcanics and is evidenced by the numerous scarps and offsets displayed within various cones and sheetflows (Beanland and Clark, 1994). Faulting also is featured in the extensive alluvial fans flanking both the Sierra Nevada and White-Inyo Mountains along the margins of the Owens Valley and in Quaternary glacial deposits along the western edge of the valley. Quaternary fluvial, lacustrine, and aeolian deposits of the basin floor also display evidence of active faulting (Beanland and Clark, 1994). Extensive active alluvial fans and uplifted and incised fanglomerates mark the range fronts of the Sierra Nevada and the White-Inyo Mountains along the margins of the Owens Valley. They record a series of progressive sedimentary and tectonic events from initiation of range-front faulting to the present (Lueddecke, et al., 1998). Pliocene lacustrine deposits throughout the basin

2 document uplift, tectonic tilting, and climate change during the last ~3 Ma. Paleo-lake beds deposited in the basin are presently tilted and exposed at elevation along the White- Inyo Mountain rangefront (Waucoba lakebeds, Bachman, 1978). This phenomenon suggests progressive subsidence, active faulting, and incorporation of the younger basin stratigraphy into the uplifted footwall of the White-Inyo block in the process of mountain building. Widespread Quaternary glacial deposits exist along the western margin of the Owens Valley, resulting from repeated glaciation of the Sierra Nevada Mountains. Despite their relative youth, these glacial deposits are prominently faulted in areas of ongoing deformation along the Sierra Nevada fault zone (Berry, 1997).

Fluvial Deposition Quaternary alluvial deposits, fluvial deposits and terraces from the Owens River and its tributaries cover much of the Owens Valley floor. The path and terrace geometry of the Owens River document tectonic basin evolution and Quaternary volcanic events (Pinter and Keller, 1990; 1992). Episodic subsidence of the Owens Valley basin has resulted in partial burial of many lava flows, glacial tills, and alluvial deposits greater than 500 ka in age (Gillespie, 1991) and much of the present basin surface is not older than ~10 k.y. (Beanland and Clark, 1994).

Pre-Cenozoic Tectonic Activity: The Inyo Thrust Fault Pre-Cenozoic tectonic activity in southeastern California has resulted in the present juxtaposition of the Precambrian to Late Ordovician rocks above mainly Mississippian-Pennsylvanian strata along the Inyo Thrust fault (Fig. 3; Stevens and Olson, 1972). A major contractional event within a Triassic tectonic belt, extending from southeastern California to Idaho, resulted in >~30 km eastward displacement of and east- vergent folding in Precambrian to Late Ordovician rocks in a 1300 sq. km block now comprising the northern Inyo and southern White Mountains (Burchfiel et al., 1970; Stevens and Olson, 1972). Movement on the Inyo Thrust fault ceased prior to the Early Jurassic, and therefore subsequently emplaced plutons (<~200 Ma) associated with the Sierra Nevada batholith are not offset by this fault (Stevens and Olson, 1972). Paleozoic rocks in the hanging wall of the Inyo Thrust fault now forming the northern Inyo and

3 southern White Mountains form a structurally coherent, large antiform entirely underlain by the Inyo Thrust fault surface and the Mississippian-Pennsylvanian rocks in the footwall. The Inyo Thrust fault is exposed along both the east side of the White-Inyo Mountains, where it is known as the Last Chance thrust, and on the west side of the range in the Owens Valley where the Cambrian and Ordovician rocks overlie the Mississippian strata along 8 km of an almost continuous exposure just east of the Tinemaha Reservoir (Fig. 3). One isolated exposure of the Mississippian-Pennsylvanian strata occurs within the Owens Valley in the Poverty Hills bedrock outcrop (Hoylman, 1974), where these units and part of an intrusive granodiorite sit several kilometers west of the White-Inyo rangefront.

The Late Cenozoic Eastern California Shear Zone Recent geodetic surveys in eastern California have shown that the Sierra Nevada Mountains and the Great Valley are moving as a rigid crustal block towards the northwest with respect to the fixed Colorado Plateau at a rate of 13 to 14 mm/yr (Dixon et al., 2000). Of this motion, 11±1 mm/yr is accommodated along the Eastern California shear zone (Fig. 4; Reheis and Dixon, 1996; Dixon et al., 2000) between the Sierra Nevada Mountains in the west and the Great Basin in the east. The Eastern California shear zone characterizes a zone of right-lateral distributed shear along NNW oriented, sub-parallel strike-slip faults that are connected by conjugate NE-striking extensional faults (Dokka and Travis, 1990). These normal faults have evolved in response to oblique transtensional stress, facilitating slip transfer within a major transcurrent zone (Reheis and Dixon, 1996). In the southern portion of this distributed shear zone, the translation of the Sierra block has been taken up by several major en-echelon fault systems concentrated near the Sierra Nevada-Basin and Range transition that include the Death Valley-Furnace Creek, the Hunter Mountain-Panamint Valley, the Fish Lake Valley, and the Owens Valley-White Mountain fault zones (Fig. 4). The Eastern California shear zone encompassing these faults defines a broad transtensional continental rift zone evolving as an incipient plate boundary in the wake of the NW-moving Sierra Nevada micro-plate. Basaltic to felsic volcanism, faulting, uplift and subsidence, and rapid river incision are the dominant geological processes currently operating within this oblique rift

4 zone and are part of the establishment of an active, intra-continental plate boundary near the western margin of North America.

Tectonics of the Eastern California Shear Zone

Plate kinematics and tectonic relations between several discrete geologic regions have determined the nature and orientation of the Eastern California shear zone. Extension in the Basin and Range Province continued on an E-W trajectory until ~10 Ma, when a change in the absolute motion of the Pacific plate initiated NNW translation of the Sierra Nevada microplate (Fig. 5; Wernicke and Snow, 1998; Brady, et al., 2000). Subduction-related Jurassic to Late Cretaceous plutonism in central and southern California produced the Sierra Nevada batholith, which was subsequently episodically uplifted and eroded, most recently beginning in the Late Miocene or Pliocene (Dalrymple, 1964; Christensen, 1966; Unruh, 1991; Wakabayashi, 2001) concurrent with the latest stages of Basin and Range extension. Subduction of the Pacific-Farallon ridge initiated the San Andreas dextral transform plate boundary ~22 Ma (Atwater, 1970; Nitchman, et al., 1990). The change in the direction of Pacific Plate motion from easterly to NW apparently prompted the separation of the Sierran microplate along a similar trajectory. Partitioning and eastward migration of associated margin-parallel deformation across the Sierran block and into the western Basin and Range initiated a broad zone of dextral transtension that includes faults of the Walker Lane and the Eastern California shear zone (Dilles, 1994; Dilles and Gans, 1995, Applegate, 1995). Present day deformation in the Basin and Range Province and the motion of the Sierra Nevada block are attributed to the onset of distributed NW-trending right-lateral simple shear, which has produced the observed pattern of transtensional fault systems (i.e., Death Valley- Furnace Creek fault zone, Owens Valley fault, Fish Lake Valley fault zone, etc.) and associated conjugate normal faults (Fig. 4; Lueddecke, et al., 1998; Nitchman, et al., 1990; Applegate, 1995).

The Owens Valley Fault A prominent segment of the Eastern California shear zone is represented by the Owens Valley fault, a 120-km-long, 3-km-wide zone of active crustal deformation within

5 this transtensional stress regime (Fig. 1, 6). The Owens Valley fault strikes ~N17W, and dips along strike range from 50° to vertical (Knopf, 1918; Lubetkin, 1980), with most typically 60°-70° (Beanland and Clark, 1994). Step-overs and fault splays along strike of the Owens Valley fault produce locally compressional and extensional stresses demonstrated by localized uplifts and pull-aparts (Beanland and Clark, 1994). The Owens Valley fault accommodates 3±2 mm/yr of dextral strike-slip motion associated with the kinematics of the Eastern California shear zone (Dixon, et al., 2000). Dextral slip is transferred from the Owens Valley fault northeastward to the dextral Fish Lake Valley fault zone by means of the Deep Springs fault, a NE-striking down-to-the- northwest normal fault in a right-stepping wrench basin (Deep Springs Valley), whose SW end lies at the latitude of Big Pine (Fig. 6; Reheis and Dixon, 1996). During prior Basin and Range extension, the nature of fault activity across the eastern Sierra-Basin and Range transition zone was primarily E-W extensional (Fig. 5), but in response to current plate motion, an overall pattern of NW-directed oblique transtension currently characterizes fault motions in the Owens Valley and the surrounding region.

Structure of the Owens Valley Basin Within and adjacent to the Owens Valley sedimentary basin exist a variety of tectonic, geomorphologic, volcanic, bedrock, and depositional features. Major fault systems bounding the Owens Valley are the Sierra Nevada fault zone, the White Mountain fault, and the Inyo fault (not to be confused with the Inyo Thrust fault) which delineate the topographic breaks between the down-dropped Owens Valley block and the adjacent ranges (Fig. 6). Vertical motion of the White-Inyo Mountains block began as much as 12-14 mya (Stockli, 1998), but their present dramatic relief was achieved during the last ~3 Ma (DePolo, 1989; Stockli et al., 2000; Pinter and Kellar, 1995; Bachman, 1978). Within the valley block, the Owens Valley fault is the dominant system of active deformation. The Owens Valley fault splits from the Sierra Nevada fault zone in the southern Owens Valley, about 3 km south of the Alabama Hills (east of Lone Pine), and both fault systems are discrete until the latitude of Big Pine. North of Big Pine, the relationship between these two fault zones is less clear and the Sierra Nevada fault zone

6 is not as well defined. The Owens Valley fault makes a left step across the Poverty Hills near the latitude of Big Pine, and continues NNW to the southeast margin of the Coyote Warp located southwest of Bishop (Ramelli and DePolo, 1987; Martel, 1984). The Coyote Warp exists here as a diffuse transition zone or ‘ramp-like’ structure (Dawers, pers. comm., 2001), and while there are no through-going faults exhibited to connect the Owens Valley fault and the Sierra Nevada fault zone, a connection at depth between the two systems is not precluded (Bateman, 1965; Pakiser, 1964; Martel, 1984). The Owens Valley basin block is an asymmetric graben consisting of multiple bedrock benches, shallowly buried (300 m or less) west of the Owens Valley fault, and in places exposed as foothills of the Sierra Nevada Mountains (Fig. 7; Gillespie, 1991). The deepest part of the basin is on the east side of the Owens Valley fault, and geophysical data indicate a sediment depth to the basement of more than 2 km (Pakiser et al., 1964). The Owens Valley fault divides the Owens Valley along its axis, separating the western bench from the deeper grabens to the east (Gillespie, 1991; Beanland and Clark, 1994). Structural discontinuity also exists along the length of the Owens Valley basin block. The deep narrow region on the east side of Owens Valley is separated by a region of shallowly buried basement in the area of Poverty Hills (Gillespie, 1991). The northern graben extends from Crater Mountain to near Bishop and the southern graben extends to the Owens Lake playa in the south.

Poverty Hills Poverty Hills is a bedrock structure located in central Owens Valley just south of Big Pine (Fig. 1). It lies directly along the main trace of the Owens Valley Fault at a 3 km left step-over of the fault and is composed of Paleozoic and Mesozoic basement rocks. Surface exposure of the bedrock in the Poverty Hills is ~13 square km, rising ~340 m above the basin floor. The Poverty Hills are separated geographically from the Sierra Nevada front by a distance of 8-10 km, most of which is covered by recent alluvial, volcanic, fluvial and glacial deposits. Poverty Hills is located centrally to a series of several large and small cinder cones (Fig. 8), collectively known as the Big Pine volcanic field. Although current fault displacement on the Owens Valley Fault is predominantly strike-slip, gravity and seismic data (Pakiser et al., 1964) indicate 1-2 km

7 of down-to-the-east vertical bedrock displacement on the fault. Whereas basement rock is deeply buried east of the Owens Valley fault, it either crops out or is shallowly buried by recent sediments on the west, except at the latitude of Poverty Hills where bedrock occurs at the surface and is shallowly buried across the width of Owens Valley.

Proposed Origins of the Poverty Hills Several mechanisms have been proposed for the occurrence of the Poverty Hills basement structure within the Owens Valley. Pakiser et al. (1964) postulated that Poverty Hills was a discrete block that had descended under the force of gravity from some altitude in the Sierra Nevada and that it came to rest in its present location as a mega-landslide. He inferred this from gravity data, by which he also interpreted the main trace of the Owens Valley fault as running on a linear trend directly beneath Poverty Hills with no step-over. Most recently, Bishop (1998; 2000) indicated that Poverty Hills may represent a deposit from a “long-runout avalanche,” one of several in Owens Valley the author cites, that traveled several kilometers out into the valley from the bounding ranges on either side. This explanation seems improbable for several reasons. Projecting the Owens Valley fault beneath the Poverty Hills structure is structurally incompatible with existing faults to the southeast, east and northwest as mapped by Martel (1984; 1989; Martel, et al., 1987), and Bateman (1965). Additionally, subsequent activity on the Owens Valley fault should have disrupted the proposed Poverty Hills slide block and in turn disclosed the trace of the underlying Owens Valley fault, however, there are no mapped through-going faults within Poverty Hills to support either this relationship, or the proposed geometry of the underlying Owens Valley fault (Bateman, 1965). The existence of the 1 Ma rhyolite dome (Moore and Dodge, 1980) just west of Poverty Hills gives a minimum age for Poverty Hills as a gravity slide, since it must predate the dome if it slid to its present locale. It is unlikely that a slide mass of these dimensions sitting directly atop the Owens Valley fault would be undisturbed given the fault activity in the last 1 Ma. Bateman (1965) states that the emplacement of Poverty Hills is due to uplifting of the crystalline basement beneath the valley fill. Martel (1989) expanded Bateman’s interpretation to state that Poverty Hills is likely a push-up structure produced by a

8 restraining left bend along the main trace of the dextral Owens Valley fault. His work on the Fish Springs fault and other faults surrounding the Poverty Hills supports this interpretation and remains compatible with Pakiser’s geophysical data (1964). Step- overs and bends present along the main trace of the Owens Valley fault are common to strike-slip fault zones, and in dextral shear zones specifically, left step-overs are structurally areas of compression and sites of potential uplift (Sylvester, 1988; Bateman, 1965; Martel, 1984, 1989; Martel, et al., 1987). Although the structural geology surrounding Poverty Hills is consistent with its location at a left step in a right-lateral oblique shear zone (Sylvester, 1988), the nature of its true genesis and geochronologic evolution have not been concluded.

9 CHAPTER 2: Field Investigation

Location, topography, methods The Poverty Hills are located at 37° 03′ 00″ N latitude, 118° 15′ 00″ W longitude, 11 km south of the town of Big Pine, and straddle the boundary between the Fish Springs and Tinemaha Reservoir USGS 7.5 minute topographic quadrangles, Inyo County, CA (Fig. 1). They rise to a maximum elevation of just over 1500 m, stand ~340 m above the Owens Valley floor, and comprise a discrete elliptically-shaped bedrock body surrounded entirely by basin fill and volcanic deposits (Fig. 8, 9). Surficial bedrock geology has been mapped by Bateman (1965) and Nelson (1966), and more recently by Hoylman (1974); structural mapping was performed during the summer field season, 2001, as part of this study (Fig. 10). Basemapping was done on relevant portions of the Fish Springs and Tinemaha Reservoir USGS 7.5 minute topographic quadrangles, enlarged from 1:24000 to 1:6000 scale, and supplemented by air photos supplied by the CA Bureau of Land Management, portions of a Landsat 7 satellite image (panchromatic band 8, 15 m resolution), and geologic quandrangles by Bateman (1965) and Nelson (1966). Predominantly structural mapping was performed, to document the presence and nature of faults, shear zones, and other structural and deformational elements, with special attention paid to stratigraphic contacts and geochronologic relationships between lithologic units. Field samples of igneous and volcanic rocks were collected for the purpose of 40Ar/39Ar isotopic age analysis.

Rock Descriptions and Stratigraphic Relationships The generalized geology of the Poverty Hills is that of an elevated bedrock region encompassed by lower-lying basin fill sediments and volcanic deposits. Within the Poverty Hills structure, two main rock types predominate: the late Paleozoic metasediments that comprise the western one-third of the structure, and the Mesozoic plutonic rock which makes up the remaining two-thirds (Fig. 10). The Paleozoic rock types have been subdivided and correlated with units in the adjacent White-Inyo Mountains (Hoylman, 1974) and the igneous rock with a unit of the composite Sierra

10 Nevada batholith (Bateman, 1965). Although the metasediments have been alternately described as likely being Sierran roof pendant rocks (Knopf, 1918; Bateman, 1965; Pakiser, 1964, Bishop, 2000), definitive correlation was done by Hoylman (1974) using microscopic petrographic, fossil, and stratigraphic evidence which showed them to be Mississippian and Pennsylvanian age units found in the White-Inyo Mountains. Explicitly detailed rock descriptions can be found in Hoylman’s work (Hoylman, 1974), therefore summary descriptions are presented here.

Metasediments The Late Paleozoic units are marine sedimentary rocks that have been highly folded prior to their intrusion by the plutonic rock and have undergone contact metamorphism proximal to the intrusive contact. The oldest units found in the Poverty Hills are members of the Late Mississippian Rest Spring Shale and include foraminifera- and crinoid-bearing fine-grained calcareous siltstone and sandstone, silty limestone, and noncalcareous clay shale. The Rest Spring Shale is exposed in a N-S trending linear band along the western-most margin of the Poverty Hills (Fig. 10) and is faulted at its stratigraphic base. These are conformably overlain by rocks of the Pennsylvanian Keeler Canyon Formation which are predominantly fusilinid-bearing fine-grained silty limestone, calcareous siltstone and sandstone, and are distinguished by a chert-nodule bearing limestone marker bed at the base. The Keeler Canyon rocks are exposed in a subparallel band east of the Rest Spring Shale outcrops and in two isolated hills in the southwestern part of the Poverty Hills that are separated by granodiorite and alluvial deposits. These units collectively total ~1130 m of Paleozoic stratigraphy (Hoylman, 1974) that has been subsequently faulted after folding and are the only outcrop of rocks of this age within the Owens Valley west of the White-Inyo rangefront. In the White-Inyo Mountains, the Mississippian-Pennsylvanian strata extend beneath the ranges and acted as the basal surface (i.e., footwall surface) of the Middle Triassic Inyo Thrust Fault (Stevens and Olson, 1972), which allowed the eastward tectonic transport of the overlying Precambrian-Ordovician strata during compressional tectonism (Fig. 3). The occurrence of the late Paleozoic strata in the Poverty Hills suggests that they may represent a remnant of the Inyo Thrust Fault footwall surface,

11 similar to other tectonic windows found on the western White-Inyo rangefront and at Jackass Flats, located east of the White-Inyo crest (Fig. 3; Stevens and Olson, 1972). The thickness and geographic location of the Poverty Hills rocks are anomalous within the known extent and thickness of Paleozoic facies in this region. The true stratigraphic thickness of the Keeler Canyon Formation in the Poverty Hills cannot be determined since the top of the Formation has been truncated by intrusion of the granodiorite pluton. However, the thickness of the limestone portion of the Rest Spring Shale exposed at this location is much greater than that of the limestone facies of this unit found in the White- Inyo Mountains due east of the Poverty Hills. From northwest to southeast, the characteristic thick shale of the Rest Spring Shale thins as the shale facies is replaced by entirely carbonate Mississippian strata. At the latitude of the Poverty Hills, the Mississippian rocks in the White-Inyo Mountains are predominantly the thick shale that the Rest Spring Shale is named for, and the limestone facies progressively thickens to the southeast. The presence of thick carbonate facies in the Mississippian rocks of the Poverty Hills can be explained by the presence of a discrete, isolated thinning event in the shale, inconsistent with the known regional pattern of this unit. The other explanation would require the right-lateral transport of these units from a location to the southeast where carbonate facies dominate, northwestward along the Owens Valley to their present latitude at the Poverty Hills.

Granodiorite The Mesozoic intrusive rock of the eastern Poverty Hills (Fig. 10) is a medium-grained equigranular granodiorite composed of clear to light grey quartz, black hornblende and biotite, and feldspar that is predominantly pink and green, or grayish. It has long been considered to be an outlier of the Tinemaha Granodiorite of Bateman (1965), the type locality of which is exposed along Tinemaha Creek in the Sierran rangefront and foothills to the west of Poverty Hills. The Tinemaha Granodiorite is Middle Jurassic in age, as determined by K-Ar analysis done by Kistler et al. (1965) which gave an age of 150-183 Ma, but the Poverty Hills exposure has not been dated. The results of Ar40/Ar39 analyses of the Poverty Hills granodiorite still in progress will be reported subsequent to this study.

12 Small pockets and strands of aplitic rock outcropping throughout the Poverty Hills granodiorite are interpreted as late-stage cooling residues. Secondary epidote has given the granodiorite its’ greenish color and is prevalent along faults where it forms well-preserved slickenlines (Fig. 11). The granodiorite is predominantly undeformed, but displays localized foliation by means of planar preferred orientation of mafic minerals and sigmoidal feldspar/quartz proximal to faults and shear zones, where it becomes progressively mylonitized such that individual mineral grains are no longer identifiable.

Fluvial/Lacustrine All other rock types in the Poverty Hills are surficial deposits, limited in size and distribution, and Quaternary in age. These include volcanic basalt flows and pyroclastics, fluvial/lacustrine deposits, and alluvium. The Owens Valley surrounding the Poverty Hills is filled to its present level with alluvium derived from the adjacent ranges. Pleistocene and Quaternary deposits of this material, including glacial till and outwash, cover the surface of the valley and are intercalated with basalts and cinder cones of the Big Pine Volcanic Field (Gillespie et al., 1983). Quaternary alluvium in the Poverty Hills is restricted to depressions in the bedrock found mainly in the north-central, central, and southwestern portions of the structure (Fig. 10), and is composed of loosely consolidated coarse pebble- to sand/silt-sized deposits derived from adjacent bedrock. These are differentiated from colluvial deposits consisting of very coarse, moderately consolidated aggregates of angular granodiorite clasts in a grus-sized granodiorite matrix. The colluvium mantles the granodioritic bedrock, especially in the northeastern exposures of the Poverty Hills, but also consists of metasedimentary clasts and matrix where it is found near these bedrock types. Similar brecciated deposits near faults and shear zones are compositionally indistinguishable from the more weathered and extensive colluvium blanketing the northeastern region. The angular bedrock clasts composing the fault breccias and colluvium are also found in hydrothermal breccias that have pervaded along many faults, shear zones, joints and fractures in the Poverty Hills.

Hydrothermal breccias have a CaCO3 matrix and also contain basalt and cinders where proximal to bedrock-volcanic contacts.

13 Tinemaha Creek is a tributary of the Owens River, flowing down from the Sierra Nevada Mountains directly west of Poverty Hills, and is fed by snowmelt from the range and fault-controlled springs as it flows to the east across the Owens Valley (Fig. 9). Upon reaching the Poverty Hills, the Creek makes a 90° turn to the north and flows around the western and northern margins where it is joined by Birch Creek before eventually feeding into the Tinemaha Reservoir on the east side of the Poverty Hills. Tinemaha Creek fluvial deposits are found on the westward- and northwestward-most margins of the Poverty Hills from the current base level of the Creek to a height of up to ~10-15 m above Tinemaha Road, and a partial vertical section is exposed in the associated roadcut. The deposits are predominantly coarse, poorly sorted channel deposits displaying repeated fining-upward sequences resembling rough, large-scale truncation and cross-bedding. Clasts are entirely derived from adjacent metasedimentary and granitic bedrock and include jagged angular pieces of walnut-to-boulder sized granodiorite and metasiltstone in a coarse sand-to-walnut sized matrix of the same composition. The fluvial deposits are interbedded in places with unsorted, unbedded colluvial deposits also composed of angular bedrock clasts in coarse pebbly matrix (Fig. 12). These deposits grade back and forth with shear zones in bedrock and fault breccia and mega-breccia deposits also exposed along the roadcut, with bedrock, colluvium and breccia exposed upslope above the vertical extent of the fluvial deposits. Lacustrine and fluvial deposits are found on the easternmost side of the Poverty Hills resting on a tilted basement block adjacent to Tinemaha Reservoir and exposed in a railroad cut on the east margin of the block (Fig. 13). The tilted fluvial/lacustrine sediments are underlain by granodiorite bedrock and are interbedded with an overlying basalt flow. All units in the sequence strike N35-55W, and dip 20-45SW, and included is a well-known and previously mined sequence of diatomite beds bounded by ash beds. The formation of these strata and the other associated lacustrine units may have occurred as a result of the damming of the Owens River by coalescing basalt flows further south from the present-day Tinemaha Reservoir, the eruptive centers of which may also have supplied the ash creating favorable conditions conducive to diatom growth (Cleveland, 1958). The stratigraphy and diatom assemblages have been meticulously described by Cleveland (1958) who interpreted the oldest sediments exposed in the railroad cut as late

14 Pliocene-earliest Pleistocene in age based on the diatom species. The uppermost fluvial/lacustrine units are overlain by very coarse colluvium composed of angular granodiorite clasts in a grus matrix. The stratigraphy exposed in the railroad cut also includes the vent region of the interbedded and subsequently tilted basalt flow, which has truncated the adjacent sedimentary beds and in places is directly in contact with the granodiorite basement of the block. The vent region has a highly altered and oxidized zone around the periphery that is interpreted as a baked margin. The vent is exposed at the southern end of the railroad cut which is vertical and strikes N-S, and it was not possible to discern the true orientation of the vent or of lava projection through the vent from the apparent cross- section produced by the cut. Its location, however, falls roughly along strike of a proposed normal fault that has been mapped by previous researchers in this area which strikes ~N65W and dips 80NE (down-to-the-NE), suggesting the fault may have provided a conduit for the basaltic magma. It is also possible that the fault post-dates the basalt, exploiting the contrast in lithologies. It is likely that this basalt was erupted into the water of the lake body, hence its direct stratigraphic succession within the lakebed sequence and supported by the appearance of subdued pillow structures and smooth, shiny pitted surfaces on the upper exposure of the flow. Also observable in the railroad cut is an angular unconformity separating steeper dipping (40-45SW) older fluvial/lacustrine and coarse colluvial beds below from gently dipping (~20SW) finer-grained fluvial/lacustrine deposits above (Fig. 13). The location of the erosional surface that created the unconformity is enhanced by a thin horizon of loose mostly football-sized basalt boulders mixed with baseball- and smaller-sized granodiorite clasts that forms a lag deposit marking the base of the unconformity. Granodioritic material is also interbedded with the upper fluvial/lacustrine units of the older tilted sequence beneath the unconformity. The basalt boulders at the unconformity can be traced laterally to the south into the predominantly overlying basalt flow and appear to be derived from it. The erosion of the basalt across the surface of the angular unconformity and subsequent tilting of the lakebeds overlying it suggests a hiatus between periods of uplift. The Plio-Pleistocene age of the oldest steeply tilted beds and the age of the gently tilted basalt (640±50 ka; Bierman et al., 1991) require that the hiatus

15 was likely brief in duration. Even younger lakebeds overlying the capping basalt were present during Cleveland’s study (1958) but these have been mostly stripped away during subsequent mining excavations. They remain preserved in the down-thrown block of the previously mapped normal fault and are exposed in the railroad cut as mentioned.

Volcanics The cinder cones and basalt flows of the Big Pine Volcanic Field surround the Poverty Hills structure and include the tilted basalt overlying the lakebeds near Tinemaha Reservoir and several other eruptive centers and flow remnants within the Poverty Hills. The ages of adjacent cones and flows range from ~1.2 m.y. to as young as ~25 k.y. found primarily by K-Ar dating (and a few Ar40/Ar39 analyses) of basalts from individual eruptive centers performed by numerous investigators spanning the last 50+ years. A compilation of this information was done in the context of this study (Fig. 14 and Table 1), but includes only one date within the Poverty Hills, that of the tilted basalt on the east near the Reservoir (640±50 ka; Bierman et al., 1991). No other volcanics within the Poverty Hills have previously been dated, but the Ar40/Ar39 analysis of three tectonically involved samples (two basalts and a basalt tephra) still in progress will be reported subsequent to this study. The rocks of the Big Pine Volcanic Field are alkali-olivine basalts with visible olivine crystals up to 2 cm and pyroxene crystals 3-5 mm in a finer-grained matrix of euhedral plagioclase laths, olivine, and pyroxene with lesser amounts of other associated mafic minerals. Most exposures are fresh and unweathered, appearing grey-black or blue-grey on broken surfaces. Xenoliths are common and include granitic inclusions derived from the underlying Mesozoic plutonic rocks through which many basalts have erupted (Gillespie et al., 1983; 1984), and lithospheric mantle xenoliths composed predominantly of orthopyroxene in a matrix of clinopyroxene, olivine, and spinel (Beard and Glazner, 1995; Ormerod et al., 1991). Dating of Big Pine basalts has been done by the traditional K-Ar method and more recently by Ar40/Ar39 analysis of matrix plagioclase laths and K-spar crystals from granitic xenoliths (Gillespie et al., 1983; 1984; Martel et al., 1987).

16 The Crater Mountain cinder cone lies NNW of the Poverty Hills and is 110 k.y. old (Fig. 14; Stone, 1993). Between Crater Mountain and Poverty Hills is the Fish Springs cinder cone (314 k.y.; Martel et al., 1987), a small cone which has been down- dropped on a normal fault and partially buried. Due west of Poverty Hills is a singular rhyolite dome, also partially buried, which is ~1.0 m.y. old (Cox et al., 1963). Southwest of Poverty Hills is Red Mountain (~240±60 k.y. old; Bishop, 2000) which flowed eastward until encountering the Poverty Hills where the flow was diverted southward around the southern margin. Southeast of the Poverty Hills at the base of the White-Inyo rangefront are three eruptive centers collectively referred to as the Split Mountain cones which range in age from 1.2 m.y. to 225 k.y. old. The 225 k.y. old sheet flow here is thought to have coalesced with the Red Mountain flow, damming the Owens River for a time (Cleveland, 1958; Gillespie, 1991), but the young ages of these flows require previous damming events to have occurred if the older lacustrine units (Plio-Pleistocene) on the eastern tilted block of the Poverty Hills were produced by such a situation. Further south and southwest of the Poverty Hills, numerous eruptions along the base of the Sierran rangefront have been faulted and partially buried. These volcanics collectively range in age from 1.2 m.y. to 25 k.y. with the oldest units existing mostly as remnants in some places interbedded with glacial deposits, and overlain by younger flows, alluvial fans, and fluvial deposits. The Big Pine Volcanic Field has been interpreted as a spatially and temporally discrete episode of fault-controlled volcanism. The structural and kinematic evolution of these faults, their temporal relationship to volcanism, local and regional stress regimes, and Owens Valley basin evolution have not been investigated but are addressed by this study. The fault conduits through which the Big Pine basalts were erupted and the Poverty Hills structure at the center of the volcanic field lie at a complex major structural intersection between the left stepover of the Owens Valley fault, the SW end of the Deep Springs fault, and the southern end of the White Mountain fault (Fig. 6). It is likely that the relationships between these structural entities and their kinematic evolution through time have a bearing on the development of the Big Pine volcanics. Within the Poverty Hills, vents and remnants of flows are preserved on the NNW, the SSW, the SE, and the already-mentioned eastern margins of the structure, as well as

17 one small flow remnant in the south-central region (Fig. 10). All have been offset, shifted, or otherwise disturbed by syn- or post-eruptive deformation of underlying Poverty Hills bedrock, and in some areas have been tectonically intercalated with slivers of bedrock. All volcanics are in direct adjacent or subjacent contact with bedrock, with flows lying directly on bedrock surfaces (except on the east margin where the flow is interbedded with lacustrine units) and vents may have been located originally along faults in bedrock.

Structure

Structural elements mapped in the Poverty Hills and used to interpret its origin include faults, shear zones, slickenlines, foliations and lineations, folds, and post-eruptive tilting/disturbance and tectonic intercalation of volcanics (Fig. 10). Bivergent imbricate thrust faults and shear zones strike predominantly NW, dipping westward in the east and eastward in the west. N-S striking normal faults within the Poverty Hills structure form a centrally located graben. A major normal fault bounds the NW margin of the Poverty Hills, down-dropping Quaternary fluvial and basin sediments to the west, and exposing deformed basement rocks in the revealed footwall (Fig. 15). This fault is similar in orientation and location to other existing previously mapped normal fault zones NW and SE of the Poverty Hills within locally transtensional regimes around the NW and SE margins of the Poverty Hills (Fig. 8). Basalt flows in the NNW, E, and SSW have been uplifted and tilted post- or syn- eruptively by underlying basement deformation (Fig. 10). Tilted vesicles, flow horizons, and columnar joints indicate some flows were likely horizontal when they cooled. Cinder cones in the S and SSW are disturbed and shifted by adjacent bedrock movement. These structures and stratigraphic relations collectively represent a positive flower structure at a left stepover along the Owens Valley Fault (Fig. 10, 16), produced within a locally transpressional structural regime superimposed on the regionally transtensional regime of the Eastern California Shear Zone (Fig. 17). Several indicators may provide general constraint of the timing and duration of uplift in the Poverty Hills. The tectonically involved basalt capping the eastern-most fault block of the Poverty Hills dates to 640±50 ka (Bierman et al., 1991) indicating active uplift during the Pleistocene in the Poverty Hills both prior to eruption of and

18 subsequent to cooling of this basalt. The oldest fluvio-lacustrine beds exposed in this block are estimated to be late Pliocene- earliest Pleistocene in age and are interbedded with granodioritic material in their upper sequence, suggesting the onset of uplift occurred near this time. Ages of other tectonically involved volcanic rocks may provide a range by which to estimate the duration of uplift. A stratified basalt tephra found on the western margin of the Poverty Hills dips gently to the east and is underlain and overlain by coarse loosely consolidated brecciated granodiorite associated with the west-bounding normal fault. The age of this syntectonically deposited tephra (in process) will also aid in constraining the timing of uplift in the Poverty Hills. The 240±60 ka (Bishop, 2000) Red Mountain basalt flow gives a minimum age for the surface expression of uplift by having been forced to flow around the positive topography (as did Tinemaha and Aberdeen Creeks) (Fig. 8). These indicators loosely constrain the age of the Poverty Hills to the latest Pliocene or early Pleistocene.

Faults and Shear Zones Faults and shear zones are pervasive throughout the Poverty Hills and present a generally decentric pattern of subparallel linear to gently arcuate traces stepping outward from the central region of the Poverty Hills to its margins and dipping back toward the center (Fig. 10). Strikes are predominantly NW and curved traces are in general concave to the center of the structure. Motion on fault planes and in shear zones is dominantly reverse-sense, as evidenced by up-dip compressive foliations and fault planes with oblique up-dip lineations and slickenlines, respectively. Dips generally steepen (to ~65° or more) inward toward the central region and shallow (to ~20°) toward the margins. Zones of deformation are discrete and narrow (0.5 to a few meters), grading rapidly to regions of undeformed bedrock that separate them (Fig.18). This bivergent pattern of compressive faulting and shearing demonstrates a complex series of imbricate thrusts descending outward from the center of the Poverty Hills that accommodate both vertical uplift and horizontal transport (Fig. 16). Back-rotation of structural slivers in the east suggests that fault surfaces at depth may resemble a poorly developed low-angle (<20°- 30°) ramp-and-flat geometry beneath the outer region of the structure, steepening to subvertical near the center where the Owens Valley fault is presumably continuous at

19 greater depth. Idealized diagrams of transpressional positive flower structures show material transported upward and outward on highly curved faults stacked one above the other (Fig. 19), but this is not entirely compatible with the observed fault geometries. In the granodiorite of the eastern half of the Poverty Hills, imbricate shear zones accommodate motion entirely within the plutonic rocks (Fig. 10), but in the northwestern and southeastern regions and where contacts with the metasedimentary units occur, faulting crosses lithologic contacts. Cataclastic fault breccias and hydrothermal breccias are commonly found at fault contacts containing clasts of either or both types of adjacent bedrock in a matrix of either crushed bedrock and/or calcium carbonate (Fig. 20). The road cut along Tinemaha Road on the north and west sides of the Poverty Hills reveals several such fault contacts where NW-directed up-dip transport has pervasively sheared the bedrock (Fig. 21). Many other excellent exposures of thrust faults occur throughout the Poverty Hills. On the western margin, an E-dipping intermediate-angle thrust with WSW sense of transport places rocks of the Rest Spring Shale above calc-schist in the footwall (Fig. 22). In the central region west of the New Era Mine, higher angle NE- directed shear zones in granodiorite display complex drag folding and rotation. Thrust faults in the southeastern Poverty Hills place rocks of the Keeler Canyon Formation above granodiorite along one strand (Fig. 23), and display stacked imbricate wedges of granodiorite exposed in a mining cut in another (Fig. 24). Normal faulting also occurs in and around the Poverty Hills. A N-S trending, steeply dipping normal fault in the central Poverty Hills bounds an elliptical sag-pond playa on the east (Fig. 25) and likely represents part of the ‘keystone graben’ seen in idealized flower structures (Fig. 16, 19). The geometry of adjacent faults indicates that of a half-graben structure beneath the playa basin and as such may be the result of trantensional motion (a ‘pull-apart graben’) (Fig. 16). A larger NNW-trending down-to- the-west normal fault extends from east of the Fish Springs cinder cone into the northern Poverty Hills, and a parallel strand may bound the adjacent alluvium-filled depression on the west (Fig. 10). Other N-S trending normal faults mapped previously by others occur immediately to the northwest (the Fish Springs fault and the unnamed rhyolite dome fault) and southeast of the Poverty Hills structure (Fig. 8) corresponding to zones of extension that occur obliquely to transpressional step-overs. A major normal fault

20 bounds the NW margin of the Poverty Hills, down-dropping fluvial deposits of Tinemaha Creek and exposing deformed and faulted Poverty Hills bedrock along with interbedded fluvial and colluvial deposits to a height of ~10-15m above the present creek (Fig. 12). The roadcut exposes bedded fluvial creek deposits above and interbedded with bedrock colluvial and breccia deposits, suggesting synchronous uplift and base level drop. The scarp of this large fault is prominent especially when viewed from a short distance (i.e., ~1km; Fig. 26), and at one point where it intersects the cut along Tinemaha Road, it can be seen to truncate thrust faulted metasedimentary and granodioritic rocks (Fig. 15).

Tectonically involved basalts All of the basalts and cinder cones in the Poverty Hills (except a basalt tephra associated with the normal fault on the west margin) appear to have undergone tilting or disturbance during the evolution of the Poverty Hills structure or were erupted synchronous to uplift. Tilting and disruption of flows and cinder cones is everywhere apparent, the most obvious location being the previously described tilted flow on the eastern fault block. An isolated flow remnant in the central Poverty Hills lies tilted directly on granodiorite with no apparent source or associated vent and has been offset vertically by 6-7 m across a small normal fault (Fig. 10). A cinder cone at the southern end of the Poverty Hills lies in sharp contact with adjacent Rest Spring Shale and is situated stratigraphically and topographically beneath it. Breccias in this area contain clasts of angular Rest Spring Shale and basalt cinders. No cinders or volcanic debris exist above the contact and two vent regions (one on the SSE and one on the SSW) are located on the sides of the cone well below its subdued and somewhat flattened summit. Flows extruded to the south from this cone (subsequently overrun by the Red Mountain flow) step downward to the SSW and systematically broken, tipped and tilted sheets suggest a brittle, post-eruptive response to uplift of the underlying bedrock (Fig. 10, 27). (Abandoned and incised stream channels of Aberdeen Creek across this same area also demonstrate southward migration in response to uplift in the Poverty Hills.) Another extremely disturbed cinder cone on the southeast margin is stratigraphically overlain by granodiorite and adjacent lacustrine sediments rest on portions of the granodiorite (Fig. 10). Breccias contain a mixture of granodiorite and cinders. No visible vent region

21 exists, but in a gravel pit dug on the south edge of the cone, stratigraphy of the cone itself appears inverted. Bright red vesicular cinders normally associated with the uppermost near-vent regions of a cone are found in lower stratigraphy at the very bottom of the pit, overlain by ~10m of finer-grained black basalt tephra and finally by fine-grained black lacustrine sands. The smooth spherical morphology of pyroclastics, especially on the east half of the cone, also suggests eruption proximal to a body of water. Tilted basalt flows and a tephra cone at the NNW corner of the Poverty Hills are the most complexly involved volcanics. A remnant columnar basalt flow caps the northern-most hill and is vertically separated from two lower flow remnants which step down to the northwest almost to Tinemaha Road (Fig. 28). All three remnants are tilted to varying degrees but may have once been parts of a continuous unit. At the northernmost edge of the same hill just above Tinemaha Road is another tilted columnar basalt that is roughly circular in shape. This outcrop and the one capping the hill above have been suggested to be volcanic necks that presumably erupted through the subjacent granodiorite (Hoylman, 1974). No evidence of baking or contact alteration (as seen surrounding the pipe located on the east side of the eastern fault block) was observed at the basal contacts of either outcrop where the overlying basalt rests on somewhat weathered and sheared granodiorite. The upper surfaces of both outcrops are smooth rounded tops of the basalt columns and display no obvious evidence of explosive volcanism or proximity to a vent. It is unclear whether the lower columnar basalt and the other three outcrops are of the same flow. The most intriguing aspect of the volcanics here is the location of a vent region on the steep hillside downslope and due NE of the hill-capping basalt. The flow is topographically above the vent, and immediately downslope from the vent region is bare undeformed granodiorite. Further downslope (due N), the granodiorite changes abruptly to black basalt tephra that is the start of a deposit that skirts the northern base of the hill and continues around the base of the next hill to the east. The tephra wraps around the northeastern and eastern base of this hill and continues on the southeastern side where it climbs both sides of a deep drainage gully before disappearing at the top of the gully at a relatively flat area on the high back (southern) side of the hills (Fig. 10). The eastern hill around which most of this tephra can be found is composed entirely of granodiorite and above the sharp contact with the

22 tephra there are no cinders or volcanic rocks present. One explanation that has been suggested to explain the presence of this tephra is that it represents a buttress unconformity produced by airfall from a nearby eruptive center (possibly the Fish Springs cinder cone or Crater Mtn.). This explanation does not account for the abrupt linear contact with the overlying granodiorite or the lack of any volcanic debris above the contact. The Fish Springs cinder cone (314ka; Martel et al., 1987) lies less than 1 km NNE from the tephra and the flow remnants above it to the west (Fig. 28) but it is unclear whether they are related. Just east of the Fish Springs cinder cone and between it and the northern topographic margin of the Poverty Hills are several other small basalt flow remnants whose origins are not known. Vertical motion on the Fish Springs fault (down- to-the-east) and an unnamed normal fault to the east that extends southeastward into the Poverty Hills (down-to-the-west) have downdropped a small triangular graben that includes the Fish Springs cinder cone, resulting in the partial burial of that cone and the relative preservation of the flow remnants on the eastern up-thrown side (Fig. 10). It has been suggested that these basalts, the Fish Springs cinder cone and the basalts and tephra of the northern Poverty Hills may all be derived from the Fish Springs cone. That would suggest uplift at such a rate that the tephra was overthrust by basement rock and the columnar remnants raised to their current elevations within the last 314 ky. It seems more likely that the basalts and tephra of the northern Poverty Hills predate the Fish Springs cinder cone, but at what point during the uplift of the Poverty Hills they were erupted is unknown. In-progress 40Ar/39Ar analysis of basalt from the upper columnar remnant may shed light on this relationship.

Folds Folding in the Poverty Hills is found primarily in the Paleozoic metasediments, which demonstrate complex asymmetric folds with steeply dipping limbs and variably plunging fold axes (20-50°). Folding in these units predates the relatively recent deformation in the Poverty Hills and was likely produced by drag-folding during eastward transport of the White-Inyo block by the Inyo Thrust Fault (of which these rocks composed the footwall surface) during compressional tectonism in the Triassic (Fig. 3; Stevens and Olson, 1972). Unconstrained rotation of folded units during uplift of

23 the Poverty Hills has obscured the original geometry of the folds but in the White-Inyo Mountains to the east, folds in these units are tight, laying over and verging to the east, consistent with drag folds formed during eastward block transport across a low angle detachment surface. Folding associated with uplift in the Poverty Hills is infrequently observed and is restricted to reverse/strike-slip faults on the western margin (Fig. 10). Fold axes in these areas trend orthogonal to kinematic and slip indicators and plunge gently (18-20°), and where overturned (NW margin in roadcut), verge updip in the direction of compression.

Other Evidence Suggesting Uplift Other observations provide circumstantial evidence of recent and possibly active uplift in the Poverty Hills. Glacial cobbles are scattered on the slopes of the isolated elliptical hill in the southwest region of the Poverty Hills. Vegetation, mainly big sage, grows in linear trends in many places, including across the central playa on three lines that strike N10W, N5W, and N5E. Big sage also grows in parallel NNW-trending lines near the location of the NNW-trending normal fault that extends into the Poverty Hills east of the Fish Spring cinder cone. Desert flowers and low shrubs are found in linear bands in the higher elevations of the Poverty Hills and do not grow beyond the restricted area.

24 CHAPTER 3: Structural Analysis

Faults and Foliations Lower hemisphere equal area stereonet plots of contoured fault planes and fault planes with slickenlines (Fig. 29-32) reveal a well-distributed pattern of deformation within the Poverty Hills, which varies somewhat from the presently observable map distribution of faults and shear zones. Imbricate thrust faults with a lesser strike-slip component and conjugate fault systems characterize the eastern part of the Poverty Hills. Higher angle strike-slip reverse faults and splay systems are found in the western region of the Poverty Hills. These variations in fault character and geometry across the Poverty Hills result from differential horizontal shearing produced by greater slip on the southern Owens Valley fault and lesser slip on the northern Owens Valley fault. Differential slip across the step-over of the Owens Valley fault is also suggestive of overall map-scale vertical-axis shearing and rotation (clockwise) of the Poverty Hills structure within the ongoing regional dextral translation of the Owens Valley fault. Fault planes with slickenlines demonstrate a significant component of horizontal motion and better resemble oblique strike slip faults with varying amounts of up-dip motion. Rotation of the structure and transpression within the step-over result in progressive reorientation of principle stresses through time which produce the observed distribution of deformation and the strongly strike-slip character of faults and shear zones. Equal area plots of foliations and foliations with lineations (Fig. 33-36) demonstrate similar distributions and geometries. It must be noted however, that foliation of bedrock suggests deformation of a more ductile nature that likely occurred at greater depth. Foliations in the Poverty Hills likely predate recent shallow brittle transpressional deformation and were probably reactivated and/or exploited by these most recent tectonic stresses. The origin of the foliations in the Poverty Hills is speculative, but due to the significant dextral strike-slip component of motion they demonstrate and the location of the bedrock on the immediate west side of the Owens Valley fault, it is possible that foliation occurred as a result of earlier strike-slip tectonic activity on the Owens Valley fault.

25 Folds in Metasediments Metasedimentary units in the western Poverty Hills have been subdivided into three separate regions for analysis of bedding orientations and interpretation of fold axes. These include metasediments of the west central, the southwestern, and the southern Poverty Hills (Fig. 10) and equal area stereo plots have been done for each (Fig. 37, 39, 41). Fold axes in the west central region have four general orientations, the most dominant trends of which cluster around N-S and plunge both N and S from 0 to ~20° (Fig. 38). Axes trending NNW-SSE plunge mainly NNW at ~50°, axes trending NNE- SSW plunge SSW from 0-20°, and axes that trend E plunge E at ~50°. Fold axes in the isolated hill of the southwestern Poverty Hills trend primarily SSW and plunge 20-35° (Fig. 40). Axes of folds in the southern region of metasediments are horizontal and trend NW-SE and also trend ESE and plunge 30-45° (Fig. 42). The axes of all three regions have a general orientation around NNW-SSE when considered together (Fig. 43), but the fold axes of the individual domains vary considerably indicating complex, possibly multigenerational folding events (Briggs et al., 2000; Stevens and Olson, 1972). This variation may also suggest that the individual exposures evaluated do not comprise a cohesive block of metasediments and may have been separately rotated horizontally (counterclockwise about vertical axes) producing the variation in axial trends while maintaining the general pattern of plunge angles. This would be consistent with motion expected within the splays of strike-slip/reverse faults present in the western Poverty Hills.

Tilted Fluvial/lacustrine Beds and Volcanics Bedding attitudes of tilted fluvial and lacustrine sediments and basalt flows on the easternmost fault block of the Poverty Hills (Fig. 13; east of Hwy. 395 at the Tinemaha Reservoir) have been plotted on an equal area stereonet (Fig. 44). All units generally strike NW and dip SW. Within the sequence, an angular unconformity divides more steeply dipping older beds below, with an average attitude of N38W/42SW, from more shallowly dipping younger beds above, whose average attitude is N48W/17SW, and indicates progressive episodic deformation. Restoring the younger beds to horizontal

26 reveals the orientation of the older beds (N34W/25SW) prior to the latest episode of tilting (Fig. 45). The combined tilting over both episodes totals 43° and suggests back- rotation of the block to the SW, possibly on a more shallowly dipping underlying reverse fault. Some of the total block rotation may also have occurred due to motion on a suspected normal fault that strikes N55W and has down-dropped granodioritic bedrock and colluvium in the northeastern-most Poverty Hills (Fig. 10). This fault may extend along the same strike beneath Quaternary alluvium and Highway 395, and may have separated the eastern-most block from the main body of the Poverty Hills. A basalt flow in the southernmost Poverty Hills appears to have undergone post- eruptive brittle deformation, possibly due to subjacent bedrock uplift. The columnar jointed sheet flow steps downward to the SW and each step consists of an upper ‘limb’ dipping shallowly to the NE (25-30°) with lower ‘limbs’ dipping more steeply (45-75°) (Fig. 46). Each ‘limb’ section consists of cohesive columnar jointed basalt that is tilted orthogonal to the dip of the upper surface of the section. The breaks in the sheet occur across horizontal axes consistently oriented N55-60°W, and each major step is 5-6m in height with 1-2 minor steps (1-2m in height, also down-to-the-SW) occurring across the long upper ‘limb’ between each major step. The sheet is 3-5m thick and rests directly on sandy alluvium mixed with small cinders. Restoring the basalt sheet to horizontal reveals that upper surface flow direction indicators trend originally N42W and N49W and plunge 1-2°E, indicating the basalt was extruded to the SE on a near-horizontal surface (Fig. 47). This is nearly perpendicular to the orientation of the present-day topographic steps and the axes of breaks between the ‘limbs’ of each step, supporting the assertion that the brittle deformation of the basalt flow was in response to changes in the underlying basement geometry rather than flow or cooling (i.e., syneruptive) features.

27 Chapter 4: Implications for Spatial/Temporal Evolution

Introduction The Owens Valley fault is a key component of the northern Eastern California Shear Zone which currently accommodates 11±1 mm/yr transtensional motion between the northwest-moving Sierra Nevada microplate (11.2-15 mm/yr) and stable North America (Fig. 4; Dixon et al., 2000; Hearn and Humphreys, 1998; Wernicke and Snow, 1998). At the latitude of Owens Valley, motion is carried primarily on two dextral faults, the Owens Valley fault (3-4 mm/yr) and the Fish Lake Valley fault zone (8±2 mm/yr), that are structurally connected by the Deep Springs fault (~1 mm/yr; Dixon et al., 2000; Lee et al., 2000), a conjugate normal fault that transfers slip from the Owens Valley fault to the Fish Lake Valley fault zone (Fig. 6; Dixon et al., 1995; Reheis and Dixon, 1996; Reheis and Sawyer, 1997; Dixon et al., 2000; Lee et al., 2001). Although the Late Pliocene (~3 Ma) to Recent extensional evolution of the Owens Valley basin is geomorphically well displayed (Gillespie, 1991), the role of the Owens Valley fault in its tectonic formation and the temporal effects of strike-slip tectonics on basement geometry have not been well constrained. The apparent lack of field evidence for strike slip motion on the Owens Valley fault has led to purely extensional, Basin and Range-style models of Owens Valley basin development (ex., Stockli et al., 2000; Gillespie, 1991; Bachman, 1978; Lueddecke et al., 1998). Inclusion of horizontal displacement during the rapid vertical separation of the Owens Valley basin/White-Inyo Mountains demonstrates the spatial and temporal formation of discrete sub-grabens in the northern and southern Owens Valley and the shallow bedrock anomaly dividing them (Fig. 7). Concurrent local transpressional uplift of the Poverty Hills, transtensional faulting, and associated volcanism occurring since the early Pleistocene along the Owens Valley fault (Big Pine volcanic field) suggest a change in local fault kinematics that may have resulted in a change in fault geometry. It is proposed that the Owens Valley fault may have stepped westward from its continuation with the White Mountain Fault to its present location beneath Crater Mountain (Fig. 6) during the latest Pliocene-early Pleistocene, in response to rapidly evolving transtensional fault patterns. This proposed change in geometry further resolves differential subsidence

28 of the Owens Valley basin and fits well within the developing regional transtensional fault pattern of the Eastern California Shear Zone. This time period also corresponds to the migration of overall slip transfer through Deep Springs Valley into the Fish Lake Valley fault zone to the east (Lee et al., 2001), shifting of the locus of shear from the Death Valley-Furnace Creek Fault Zone to the Owens Valley fault (Lee et al., 2001), a decrease in slip rates on the White Mountain Fault and the Death Valley-Furnace Creek fault zone, and an increase in slip rate on the Fish Lake Valley fault zone north of its intersection with the Deep Springs fault (Dixon et al., 2000; Reheis and Dixon, 1996). The recognition of the Poverty hills as a structural/tectonic feature provides important chronological and kinematic constraints for the tectonic evolution of the Owens Valley and Eastern California shear zone. The spatial, temporal, and structural genesis of the Poverty Hills integrated with a synthesis of published research leads to a structural geochronology describing Owens Valley basin formation (Table 2). Evaluation of superimposed regional and local stress regimes in the Owens Valley and the Poverty Hills (Fig. 17) allows the formulation of a spatial and temporal model of Owens Valley basin formation (Fig. 48, 49) and the evolution of the Eastern California shear zone in this region. Geometric modeling of uplift combined with calculated and recently published slip rates allows an estimate of the slip rate on the northern Owens Valley fault since the onset of deformation in the Poverty Hills (Fig. 50) and calculation of estimates of horizontal motion on the Owens Valley and White Mountain faults since the middle Pliocene (Table 3).

Changing kinematics and volcanism The Owens Valley Fault was once continuous with the White Mountain Fault (Fig. 48; Stockli, 1999; dePolo, 1989), but the White Mountain fault presently accommodates a lesser amount of strike-slip motion than the Owens Valley fault. This prior configuration requires a right stepover of the Owens Valley fault at the latitude of the Poverty Hills in order to join the White Mountain fault (Fig. 48). This contrasts with the present location of the Owens Valley fault which steps left at the Poverty Hills and continues north beneath Crater Mountain almost to Bishop, where another left step crosses the Coyote Warp to join the Round Valley and other faults of the Sierra Nevada

29 fault system (Fig. 6; Martel et al., 1987; 1989; Dawers, personal communication, 2001; 2002). The migration of the Owens Valley fault to west of Big Pine (Fig. 49) changed the geometry and kinematics of the Owens Valley fault, producing the transpressional uplift at the Poverty Hills left step-over and the associated locally transtensional fault splays, which may have provided pathways for volcanic eruptions. The Big Pine volcanic field does not necessarily constrain the initiation of the faults providing conduits for the individual eruptive centers, but the age of onset of the eruptive sequence may indicate a change in local or regional fault kinematics allowing these faults to become conduits. Eruption of similar mantle-derived basalts throughout the Basin and Range transition zone suggests that mantle conditions beneath the western North American continent are not unique to the Big Pine area, but upper crustal conditions allowing eruption of this material are (Manley et al., 2000). The concurrent timing of transpressional uplift of the Poverty Hills in the context of evolving regional transtension associated with the late Tertiary-Quaternary development of the Eastern California shear zone supports the likelihood that the eruptive and deformational events are not unrelated.

Regional spatial and temporal evolution of the Eastern California shear zone: Previous work Geologic (Dokka and Travis, 1990a) and geodetic (Dixon et al., 2000; Savage et al., 1990; Sauber et al., 1994; Gan et al., 2000; Miller et al., 2001) data indicate that the Eastern California Shear Zone accommodates 6-14 mm/yr of northwest-directed right lateral slip and comprises ~20-25% of the relative motion between North America and the Pacific plate. Shear is transferred northeastward from the Owens Valley and Fish Lake Valley fault zone systems through a complex system of ENE-trending sinistral and normal faults in the Emigrant Peak and Excelsior Mountain region to the NW-trending faults of the Walker Lane (Fig. 4; Oldow et al., 2001; Reheis and Dixon, 1996; Reheis and Sawyer, 1997). In the southern Owens Valley, a 120 km long, 20 km wide concentrated dextral shear zone extends from the south end of the 1872 Owens Valley earthquake rupture (Beanland and Clark, 1994) to the north end of the 1992 Landers earthquake surface break in the Mojave Desert (Peltzer et al., 2001). This zone is

30 continuous through the sinistral Garlock fault which shows no evidence of recent left- lateral slip and this is interpreted to represent an oscillatory strain pattern between interacting conjugate fault systems (Peltzer et al., 2001). The slip rate of the dextral shear zone below ~5km depth is 7±3mm/yr on a vertical fault plane, which is 2 to 3 times the geologic rates estimated on NW trending faults in this region (Peltzer et al., 2001). This indicates current strain accumulation in the shallow crust exceeding the geologic slip rate and suggests the possibility of a large magnitude earthquake in order to reconcile the difference in slip being carried at depth (Peltzer et al., 2001). Similarly, geodetic slip rates for the Owens Valley fault (7±2 mm/yr) and the White Mountain fault (6±2 mm/yr; Dixon et al., 2000) are much higher than geologic slip rates (2±1 mm/yr on OVF from Beanland and Clark, 1994; 1.9-3.8 mm/yr on OVF from Lee et al., 2000; 1.8 mm/yr on WMF from Schroeder et al., 2002). The difference reflects the fact that geodetic estimates integrate motion across the entire fault zone and indicate overall differential motion at depth, whereas geologic estimates reflect motion on well developed faults strands with measureable surface expression (Dixon et al., 2000). Transient slip rates and temporal kinematic variations demonstrate the inherent complexity in identifying and reconstructing past strain patterns and interactions among faults of the ECSZ.

Onset of the Eastern California Shear Zone: Local fault ages and offsets Geologic events associated with Owens Valley basin formation considered here include the evolution of basin block geometry, the uplift of the Poverty Hills, and the proposed evolutionary westward step of the Owens Valley fault. The structural and kinematic nature of these events require regional transtension and local transpression in addition to the observed large normal component of motion separating the basin from the bounding ranges. While the structural and kinematic evolution of the Owens Valley fault is not entirely known, the onset of strike-slip and conjugate normal faulting on other individual faults in the immediate vicinity are fairly well constrained. Reconstructions of Basin and Range extension and the onset of the northwestward movement of the Sierra Nevada microplate demonstrate that the Eastern California Shear Zone initiated ~8-11 Ma in the western part of the Province (Fig. 5; Wernicke and Snow,

31 1998; Brady et al., 2000; Reheis and Sawyer, 1997). Right lateral motion on the Fish Lake Valley fault zone began ~10 Ma (Late Miocene) consistent with the change in the orientation of the extensional axis from west-southwest to west-northwest and the beginning of right-lateral shear in the western Basin and Range (Wernicke and Snow, 1998; Wernicke et al., 2000; Zoback et al., 1981; Reheis and Sawyer, 1997). Extensional faulting between 6.9 and 4 Ma resulted in the opening of Fish Lake Valley and initiation of the Deep Springs fault (Reheis and Sawyer, 1997). This change in tectonism corresponded to renewed uplift of the Sierra Nevada at 6-4 Ma (Unruh, 1991), a change in the absolute motion of the Pacific plate (Harbert, 1991) and initiation of the construction of the present-day western Basin and Range topography (Nitchman et al., 1990). Uplift of the Sierra Nevada Mountains accelerated ~4.5 Ma, (associated with the northward migration of the Mendecino triple junction (Atwater and Molnar, 1973) and coinciding with increasing slip on the San Andreas Fault and the opening of the Gulf of California (Hay, 1976)). Deposition of the Waucoba lakebeds immediately preceded the major uplift of the White-Inyo Mountains relative to Owens Valley, which also began in the Pliocene (~3 Ma; Manley et al., 2000; Gillespie, 1991; Bachman, 1978; dePolo, 1989), corresponding to a period of major extension and volcanism throughout the region east of the southern Sierra Nevada. (This has been related to the proposed delamination of the lower crust and lithospheric mantle into the asthenosphere beneath eastern California (Henry and Perkins, 2001; Manley et al., 2000)). South of the Garlock Fault, faults of the Eastern California shear zone have accumulated ~65 km of dextral offset since the late Miocene, and northward, faults of the Walker Lane have accumulated 60-75 km of dextral offset in the last 10 Ma (Fig. 4; Reheis and Dixon, 1996). Between 40-50 km of northwest directed dextral offset has occurred from the central Fish Lake Valley fault zone to the northern Death Valley- Furnace Creek fault zone, with 30 km of that since the beginning of the Pliocene (Fig. 6; Reheis and Sawyer, 1997; Hamilton, 1994). The Death Valley fault zone has accumulated ~35 km of dextral slip since ~13 Ma (Butler et al., 1988) and the Hunter Mountain Fault has had 8-10 km of right slip since ~3 Ma (Fig. 6; Burchfiel et al., 1987). Several estimates of the total dextral horizontal offset on the Owens Valley-White Mountain fault zone have been made, most recently by Lee et al., (2001) who suggested

32 ~3.8-13.3 km based on Stockli’s (1999) minimum age of onset and their own local offset measurements and calculated earthquake recurrence interval. This is in general agreement with estimates by Moore and Hobson (1961) based on correlation of the Independence dike swarm, and by Ross (1962) based on an offset Cretaceous pluton, but lower than the ~20 km suggested by Beanland and Clark (1994) during their detailed study of surface ruptures produced during the 1872 earthquake, and significantly lower than the large offsets (ranging ~30-75 km) observed on corresponding major strike-slip faults in this region of the Eastern California shear zone. The Queen Valley Fault is a NE-striking, down-to-the-northwest conjugate normal fault similar to the Deep Springs Fault, forming the northern termination of the White Mountain fault (Fig. 6). The opening of the Queen Valley basin at 3.0 ± 0.5 Ma kinematically requires dextral slip on the Owens Valley-White Mountain fault zone to have been occurring (Stockli, 1999). It is relatively certain that strike-slip motion was occurring on the Owens Valley fault by ~3 Ma when the latest, most rapid period of W-I Mountain uplift began to occur (Bachman, 1978; Stockli et al., 2000; dePolo, 1989; Reheis and Sawyer, 1997; Gillespie, 1991). Stockli (1999) has suggested that the opening of the Queen Valley at 3.0 Ma may record the onset of strike-slip faulting in the Owens Valley. The ages and offsets of the adjacent strike-slip and conjugate normal faults are greater than 3 Ma (~5 Ma for the Deep Springs fault, 10 Ma for Fish Lake Valley fault zone, and up to 75 km offset) suggesting that strike-slip motion may have been occurring across Owens Valley prior to 3 Ma. The implicit nature of conjugate normal slip-transfer faults such as the Deep Springs fault suggests (by its inception at ~5 Ma) that it evolved to accommodate motion between major strike-slip faults in an evolving transtensional fault zone (Dixon et al., 1995), which in turn suggests that strike- slip motion was being carried on faults in the Owens Valley by as early as 5 Ma.

Slip rates and paths through time Presently the Fish Lake Valley fault zone accounts for at least half the present rate of Pacific-North American plate boundary shear associated with the northern Eastern California shear zone (Wernicke and Snow, 1998; Reheis and Sawyer, 1997). Changes in lateral slip rates through time (since the late Miocene) for the Fish Lake Valley fault zone

33 have been constrained by Reheis and Sawyer (1997). The lateral slip rate decreased from ~6 to ~3 mm/yr from the late Miocene to early Pleistocene (the decrease of lateral slip rate in the Pliocene on the Fish Lake Valley fault zone is due to the onset of basin- opening extension), then increased to ~11 mm/yr during the middle Pleistocene, and decreased to ~ 2-3 mm/yr during the late Pleistocene. The large increase during the middle Pleistocene is linked to the increase in vertical slip rates on the Fish Lake Valley fault zone and the Deep Springs fault at the time of the Bishop ash/Long Valley caldera eruption (0.76 Ma; Reheis and Sawyer, 1997). The marked increase in vertical slip rates after the eruption of the Bishop ash also occurred on the White Mountain fault (dePolo, 1989) and, in terms of Quaternary subsidence of Owens Valley, on all basin-bounding faults (Gillespie, 1991). Slip between the Owens Valley fault and the Fish Lake Valley fault is presently transferred through the Deep Springs fault, a NE-striking down-to-the-NW conjugate normal fault, characteristic of connecting faults between the en echelon pattern of major strike-slip faults in this region of the Eastern California shear zone (Dixon et al., 1995). The 25-km long Deep Springs Valley is bounded on the SE by the Deep Springs fault which initiated at ~5 Ma; however movement on the Deep Springs fault was not sufficient to defeat the flow of drainages from the White Mountains into Eureka Valley (NW to SE across the Deep Springs fault) until after the 0.76 Ma eruption of the Bishop ash (Reheis and Sawyer, 1997). Several north-striking normal faults step eastward from the southern end of the White Mountain fault into the Waucoba embayment and cut Pliocene rocks (Reheis and Dixon, 1996; Nelson, 1966). The southern-most of these faults are northeast-striking, falling along strike or parallel to the Deep Springs fault, and have Quaternary offset (Reheis and Dixon, 1996; Nelson, 1966). These most recently active northeast-striking faults likely demonstrate the present path of slip transfer from the Owens Valley fault near Poverty Hills to the Deep Springs fault. A line drawn on the trace of the Deep Springs Fault intersects the Owens Valley fault to the SW at the location of the Poverty Hills. Presently, the slip rate on the 80-km long Fish Lake Valley fault zone increases northward from ~3 to ~8 mm/yr, with the increase coinciding with the intersection of the Deep Springs fault (Reheis and Dixon, 1996; Reheis and Sawyer, 1997). Currently in

34 Owens Valley, the slip rate decreases by ~1.3 mm/yr from the Owens Valley fault in the south to the White Mountain fault in the north, and the Deep Springs fault has accommodated ~1.0 mm/yr of extensional slip since the middle Pleistocene (Lee et al., 2001; Reheis and Dixon, 1996; Savage et al., 1990). The vertical slip rates on the Fish Lake Valley fault zone and Deep Springs fault were higher in middle Pleistocene than before or after (Reheis and Dixon, 1996), corresponding to rapid movement on faults bounding the Owens Valley following eruption of the Long Valley caldera (Gillespie 1991). The Owens Valley fault is currently the principal locus of dextral shear for the northern Eastern California shear zone, but previously more displacement was accommodated by the Death Valley-Furnace Creek fault zone (Savage and Lisowski, 1995). Dixon, et al., (1995) have proposed a kinematic model for the Owens Valley region of the Eastern California shear zone in which the locus of right lateral shear has shifted from the Death Valley-Furnace Creek fault zone to the southern Owens Valley fault since ~1.5 Ma. This is consistent with the late Pleistocene to Holocene rates and slip distributions of Reheis and Sawyer (1997) which demonstrate the shift in slip accommodation away from the Death Valley-Furnace Creek fault zone into the southern Owens Valley-Deep Springs-Fish Lake Valley fault zone system. Based on these temporal and kinematic relationships, Lee et al., (2001) place initiation of horizontal slip transfer along the Deep Springs fault at ~1.7 Ma, capturing ~0.7 mm/yr of Owens Valley fault slip, corresponding to reduction in slip rates along the White Mountain fault and the Death Valley-Furnace Creek fault zone (Reheis and Sawyer, 1997; Reheis and Dixon, 1996). Thus, the strike-slip strain path migrated from northern Owens Valley through Deep Springs Valley ~1.7 Ma (latest Pliocene-earliest Pleistocene), as did the locus of shear from the Death Valley-Furnace Creek fault zone to the southern Owens Valley fault ~1.5 Ma, corresponding to the decrease in slip along the White Mountain Fault (Lee et al., 2001). The proposed transtensional westward-step of the Owens Valley fault (1.5-1.7 Ma) corresponds temporally to the change in overall strain path from northern Owens Valley to Deep Springs Valley and fits well within the developing pattern of transtensional deformation in this part of the Eastern California shear zone.

35 Basin Block Geometry, Subsidence, and Faulting The geometry of the Owens Valley basin block may be explained by the spatial and temporal evolution of northwest-directed transtension distributed across the Owens Valley fault and the rangefront faults resulting in a change in Owens Valley fault geometry. The Owens Valley basement is a composite of several smaller blocks that have subsided differentially between the bounding ranges. The smaller blocks step down to the east from the Sierra Nevada fault system bounding it on the west (Gillespie, 1991). The deepest portions of the basin are linear N-S trending grabens on the eastern-most side of the block (Fig. 7) between the Owens Valley fault and the Inyo Mountains fault south of the Poverty Hills, and between the Owens Valley fault and the White Mountains fault north of the Poverty Hills (Gillespie, 1991; Pakiser et al., 1964; Bateman, 1965). Changes in depth from the top of the basin fill to the underlying bedrock along the length of the valley have been derived from geophysical data (Pakiser et al., 1964; Bateman, 1965). The deep narrow graben south of the Poverty Hills is as much as 2.1 km deep, but the basin abruptly begins to shallow north of Independence and bedrock is only up to ~300 m deep at the latitude of Poverty Hills and the Tinemaha Reservoir (Pakiser et al., 1964). North of the Poverty Hills, the basin again deepens, but only to about half the depth to the bedrock (1.2 km) as it is to the south (Fig. 7; Pakiser et al., 1964). This suggests differential motion between smaller blocks making up the composite Owens Valley block during White-Inyo Mountain block uplift and Owens Valley graben formation. Geophysical data also indicate the Owens Valley fault has accommodated 1-2 km of down-to-the-east motion beneath the present basin floor primarily in the last ~3 m.y. concurrent to the rapid uplift of the White-Inyo Mountains block that bounds Owens Valley on the east (Pakiser et al., 1964; Bachman, 1978; dePolo, 1989; Gillespie, 1991). Subsidence of the southern Owens Valley began ~5-6 Ma, prior to subsidence of the northern part of the basin, and the geomorphic evidence of Matthes (1950) shows that the southern Owens Valley was ~1 km deep during the Pliocene (Gillespie, 1991). Owens Valley currently displays ~2 km of topographic difference between the adjacent ranges and the present basin surface. Combined with depths to the basin bedrock, the southern Owens Valley has subsided ~4 km in the last ~5-6 Ma with respect to the crest of the Sierra Nevada and White-Inyo Mountains. As much as half of the total subsidence

36 of the southern Owens Valley has occurred across the Owens Valley fault as compared to about one third in the northern Owens Valley. Northwest-directed transtension through time may explain the observed basin geometry. Early in Owens Valley basin evolution (5-6 Ma), the southern Owens Valley fault and the Inyo Fault would have accommodated all of the differential block motion between Owens Valley and the Inyo Mountains south of Poverty Hills (Gillespie, 1991; Matthes, 1950). Since at least 3 Ma (Stockli, 1999), regional transtensional kinematics would have resulted in the formation of a pull-apart basin (the southern graben), bounded on the north by the right step from the southern Owens Valley fault to the White Mountain fault at the location of the present-day Tinemaha Reservoir (Fig. 48). The shift in the slip path away from the White Mountain fault into the Deep Springs fault at ~1.7 Ma and the onset of transpressional uplift in the Poverty Hills suggest the inception of the west-stepping northern strand of the Owens Valley fault (Fig. 49). Vertical and horizontal motion passed to the northern Owens Valley since this time will have been distributed across both the northern Owens Valley fault strand and the White Mountain fault (minus slip captured by the Deep Springs fault), producing the observed graben between them (Fig. 49). This model requires the rapid deepening of the structurally younger northern graben (up to 1.2 km) in the last 1.7 Ma. Although this cannot be constrained, the suggestion is supported by the previously discussed extremely high slip rates (both vertical and horizontal) on adjacent faults during this period, especially following the eruption of the Long Valley caldera and the dramatic regional relief achieved across the White-Inyo rangefront since only 3 Ma. The spatial and temporal geometry of these events is consistent with the development of en echelon steps that form in normal fault arrays and the transcurrent propagation and evolution of dextral shear zones. Modeling Owens Valley basin evolution to include regional transtensional deformation of the developing northern Eastern California shear zone better resolves differential basin subsidence and basin block geometry, slip partitioning and distribution across the northern Owens Valley through time, and the initiation and onset of uplift of the Poverty Hills. This model further demonstrates kinematic and temporal fault relations at the complex major structural intersection of the northern and southern Owens Valley faults, the White

37 Mountains fault, and the Deep Springs fault occurring at the Poverty Hills, and provides insight to their recent evolution and interaction within this key region of the northern Eastern California shear zone.

Slip and uplift estimate/calculations Slip rates decrease from south to north along the Owens Valley-White Mountain fault system at the latitude of the Poverty Hills, presumably due to the capture of 15-40% of Owens Valley fault slip by the Deep Springs fault since ~1.7 Ma (Lee et al., 2001). Current slip rate estimates for the Owens Valley fault are ~3-4 mm/yr, of which 0.7 mm/yr is transferred through the Deep Springs fault and on to the Fish Lake Valley fault zone. Schroeder et al. (2002) have calculated a late Pleistocene rate of 1.8 mm/yr for the White Mountain fault. The reduction in the slip rate on the White Mountain fault is greater (0.5-1.5 mm/yr) than the amount captured by the Deep Springs fault (0.7 mm/yr) since 1.7 Ma. This suggests that an amount of slip is being transferred by an alternate route, i.e., the northern Owens Valley fault. An estimate of the slip rate for the northern Owens Valley fault would provide a better understanding of how slip is partitioned across faults north of the Poverty Hills. The age estimate for the Poverty Hills (1.5-1.7 Ma), its geometry, and the resolution of principal stresses (Fig. 17) give a constant slip rate since 1.7-1.5 Ma of 1.27-1.44 mm/yr for the northern Owens Valley fault (Fig. 50). The combined slip rates for all three faults (White Mountain fault, 1.8 mm/yr; Deep Springs fault, 0.7 mm/yr) suggest a slip rate of 3.77-3.94 mm/yr for the southern Owens Valley fault (Fig. 50). This rate falls within current slip rate estimates for the southern Owens Valley fault and accounts for slip distribution north of the Poverty Hills. Horizontal offset estimates from calculated and published slip rates (assumed constant) are shown in Table 3. The horizontal offset for the period 3.0 Ma to 1.7-1.5 Ma for the combined Owens Valley-White Mountain fault system equals 4.90-5.91 km, thus the total offset from 3 Ma to the present is 10.81 to 12.32 km, falling within estimates by Lee et al., (2001) for the same period (3.8-13.3 km).

38 Owens Valley fault offset: Implications for Poverty Hills bedrock provenance The estimate of 4.90-5.91 km horizontal displacement on the Owens Valley fault for the period 3.0 Ma to 1.5-1.7 Ma (Table 3) leads to speculation regarding the true provenance of bedrock units exposed in the Poverty Hills. As described in a previous chapter, stratigraphic thicknesses of the Paleozoic metasediments of the Poverty Hills are not consistent with the regional trends of thickening and thinning of these strata throughout southeastern CA and western NV. The dislocation of the Poverty Hills metasediments may be explained by dextral strike slip translation of bedrock along the west side of the Owens Valley fault, but the true amount of movement is unknown. Characteristics of the Poverty Hills granodiorite (included with the Tinemaha granodiorite, 150-183 Ma, by Bateman (1965) but not previously dated) also suggest that it may be relocated. The final episode of Sierra Nevada plutonism occurred in the Late Cretaceous and is divided by the 148 Ma Independence Dike Swarm, which separates the youngest plutons from older ones by their lack of Independence dikes. Although the Tinemaha is pervasively intruded by Independence dikes, none occur in the Poverty Hills granodiorite, suggesting that it may be younger than previously assigned. This in turn suggests the Poverty Hills granodiorite may not be genetically related to the Tinemaha granodiorite. Again, dextral transport allows that the Poverty Hills granodiorite may actually be a portion of another pluton yet to be identified in the White-Inyo Mountains. Restoring the rocks of the Poverty Hills to the SSE, according to the offset estimate above, places them at the latitude of the northern Santa Rita Flat pluton in the Inyo Mountains at 3 Ma (Fig. 2). A larger amount of offset (~20 km) such as was estimated by Beanland and Clark (1994), relocates the Poverty Hills units to the SW corner of the Santa Rita Flat pluton, but requires strike-slip motion to have occurred on the Owens Valley fault prior to 3 Ma (slip rates used in this study give an offset of ~20 km since ~6.7-6.9 Ma). Conversely, the age of the Santa Rita Flat Pluton (163-164 Ma, Vines and Law, 2000; Chen and Moore, 1979) and the presence of Independence dikes make it unlikely that the Poverty Hills granodiorite is from the same or a genetically related pluton. An even larger offset (31.29-33.49 km since 10 Ma, again using the above slip rates) puts the Poverty Hills rocks near the latitude of the Hunter Mountain

39 batholith (~180 Ma, Niemi et al., 1997; some Independence dikes) in the central Inyo Mountains (Fig. 2), also an unlikely correlation. Two smaller plutons adjacent to the Hunter Mountain batholith (Paiute Monument and Pat Keyes plutons) have no dates available, and it appears on geologic maps that they may or may not contain some Independence dikes, leaving the possibility that the Poverty Hills granodiorite may be related to these smaller intrusions. The varying slip rates through time characteristic of adjacent faults, their greater ages, and the large displacements which have occurred on them suggest that the true provenance of the Poverty Hills bedrock units may be a substantially further distance to the SE than 3 Ma-offset-estimates allow. In this context, these geologic indicators, within the regional evolutionary framework of the northern Eastern California shear zone, indicate that the Owens Valley fault may have a greater age and horizontal displacement than previously recognized.

40

Fig. 1. Schematic map of geologic features in the Owens Valley region. Modified from Martel et al., 1987.

41 Fig. 2. Lithologic map of Owens Valley region. From Beard and Glazner, 1995.

42 Fig. 3. The Inyo Thrust and related structures. Modified from Stevens and Olson, 1972.

43 Fig. 4. Map of late Cenozoic faults of the Eastern California Shear Zone from the Garlock Fault (GF) to the Walker Lane (WAL). Faults (F) and fault zones (FZ): WMF, White Mountains; OVF, Owens Valley; DSF, Deep Springs; EVF, Eureka Valley; SVF, Saline Valley; HMF, hunter Mountain; TPF, Towne Pass; PVF, Panamint Valley; FLVFZ, Fish Lake Valley; FCFZ, Furnace Creek; DVFZ, Death Valley. ML is Mono Lake; WL is Walker Lake; EM is Excelsior Mountains. Modified from Reheis and Dixon, 1996

44 Fig. 5. Relative motion of the Sierra Nevada with respect to the Colorado Plateau since 36 Ma. Northwesterly motion of SN started ~10 Ma. From Wernicke and Snow, 1998.

45 PH

Fig. 6. Faults of the northern Eastern California Shear Zone. Adapted from Lee et al., 2001. Faults: DVFC – Death Valley-Furnace Creek; DSF – Deep Springs; FLV – Fish Lake Valley; HMF – Hunter Mountain; IF – Independence; INF – Inyo; OVF – Owens Valley; PVF – Panamint Valley; QVF – Queen Valley; TPEF – Towne Pass-Emigrant; WMF – White Mountain.

46 Fig. 7. Map of Owens Valley with bounding ranges. Location of northern and southern grabens shown on east side of basin. From Gillespie, 1991.

47

Crater Mtn. 110 ka Owens Valley Fault Zone White Fish Mtns. Springs Owens R. Hill Fish Springs Cinder Cone 314 ka

Birch Cr.

Rhyolite Dome 1.0 Ma Tinemaha Rsvr. PH 640 ka Tinemaha Cr.

Red Mtn. 240 ka 225 ka

Split Mtn. Cones

1 km Owens Valley 460 ka Fault Zone 1.2 Ma

Fig. 8. Enhanced Landsat 7 image of Poverty Hills and adjacent geologic and geographic features. Ages given for dated volcanics of Big Pine volcanic field.

48

Fig. 9. View to the east of the Poverty Hills.

49 Fig. 10. Structural and geologic map of the Poverty Hills compiled from structural mapping and field-collected data. Taylor, 2002.

50

Fig. 11. Slickenlines on granodiorite boulder (not in place). East central Poverty Hills.

51 Fig. 12. Fluvial and colluvial deposits on NW margin of Poverty Hills, ~8-10 m above Tinemaha Creek.

52

Fig. 13. Angular unconformity in Plio-Pleistocene lacustrine sediments. Eastern fault block near Tinemaha Reservoir.

53

Fig. 14. Map of Big Pine Volcanic Field with eruption ages of dated volcanics and local faults.

54 Fig. 15. Shaft at uplifted contact between granodiorite and Paleozoic rocks. Fault breccia of W-bounding normal fault on right. View to E.

55 Fig. 16. Structural and geologic cross section A-A′ (Fig. 10), Poverty Hills.

56

Fig. 17. Schematic map of Owens Valley fault zone showing orientations of regional and local principal stresses.

57 Fig. 18. Shear zone in granodiorite. East side Poverty Hills. View to the north.

58

Transpressional Push-Ups

From Sylvester, 1988

Fig. 19 Models of transpressional push-ups, or positive flower structures. From Sylvester, 1988.

59 Fig. 20. Breccia at fault contact. Mississippian shale clasts in CaCO3 matrix. Southeastern Poverty Hills.

60 Fig. 21. Tilted Quaternary basalt flow capping deformed granodiorite on the north edge of Poverty Hills. View to the southeast.

61 N E

Fig. 22. Oblique reverse fault in Mississippian shale. SW margin of Poverty Hills. View to the northeast.

62 Fig. 23. Mississippian shale thrust above Pennsylvanian siltstone. SE side Poverty Hills. View to NW. Jake for scale.

63 Fig. 24. Stacked wedges of sheared and faulted granodiorite. SE Poverty Hills.

64 Fig. 25. Elliptical sag pond playa in structural half-graben. Central Poverty Hills. View to the W.

65

Inyo Mtns.

Poverty Hills Fault Scarp

Fig. 26. NW and W margin of Poverty Hills showing scarp of W-bounding normal fault. View to SE.

66 Fig. 27. NE- and SW-dipping ‘limbs’ of columnar basalt sheet stepping down to the SW across NW-trending basement steps. View to the SE.

67 Fish Springs Hill Crater Mtn. Fish Springs Fault

Fish Springs Cinder Cone

Basalt Flow

Fault

Fig. 28. View north from northwestern Poverty Hills.

68

Fig. 29. Contoured equal area stereographic projection of faults in granodiorite.

Fig. 30. Equal area stereographic projection of faults with slickenlines in granodiorite.

69

Fig. 31. Contoured equal area stereographic projection of faults in metasediments.

Fig. 32. Equal area stereographic projection of faults with slickenlines in metasediments.

70

Fig. 33. Contoured equal area stereographic projection of foliations in granodiorite.

Fig. 34. Equal area stereographic projection of foliations with lineations in granodiorite.

71

Fig. 35. Contoured equal area stereographic projection of foliations in metasediments.

Fig. 36. Equal area stereographic projection of foliations with lineations in metasediments.

72

Fig. 37. Equal area stereographic projection of bedding in metasediments in west central Poverty Hills.

Fig. 38. Distribution of fold axes in metasediments of the west central Poverty Hills.

73

Fig. 39. Equal area stereographic projection of bedding in metasediments of isolated hill in southwestern Poverty Hills.

Fig. 40. Distribution of fold axes in metasediments of isolated hill in southwestern Poverty Hills.

74

Fig. 41. Equal area stereographic projection of bedding in metasediments of the southern Poverty Hills.

Fig. 42. Distribution of fold axes in metasediments of the southern Poverty Hills.

75

Fig. 43. Distribution of fold axes in metasediments of all domains.

76

Fig. 44. Equal area stereographic projection of bedding of sediments and basalt flow east of Hwy. 395.

Fig. 45. Tilt restoration of younger (pink) and older (blue) fluvial/lacustrine beds of tilted fault block.

77

Fig. 46. Equal area stereographic projection of ‘limbs’ (Fig. 27) of disturbed basalt flow in southern Poverty Hills.

Fig. 47. Tilt restoration of basalt flow (blue) with directional flow lineations (magenta) in the southern Poverty Hills.

78

Fig. 48. Spatial and temporal model of NW-directed oblique transtension. Owens Valley basin prior to ~1.7 Ma.

79

Fig. 49. Spatial and temporal model of NW-directed oblique transtension. Present Owens Valley basin.

80 Fig. 50. Slip distribution with calculated (this study) and published slip rates for faults in the vicinity of the Poverty Hills and Owens Valley.

81

Table 1: Ages of dated BPVF basalts

Location Age Researcher/Source

Crater Mountain ~110k Stone, 1993

Fish Springs cinder cone 314k Martel et al., 1987

Rhyolite dome ~1my Cox et al., 1963 (in Gillespie, 1991)

Basalt of E PH fault block ~600k Cleveland, 1958 (in Gillespie, 1991)

Basalt of E PH fault block 640±50k Bierman et al., 1991 (in Bishop, 2000)

Split Mountain cones ~1my B. Turrin, pers.comm.(in Gillespie, 1991)

Split Mountain, S cone 1.2my A. Jayco, pers. comm.

Split Mountain, W cone 460k A. Jayco, pers. comm.

Split Mountain, sheet flow 225k A. Jayco, pers. comm.

Red Mountain 240±60k in Bishop, 2000, not cited

Goodale Creek 25k Stone et al., 1993

Armstrong Canyon 25k Stone et al., 1993

Sawmill Creek 119k Gillespie et al., 1984 (in Gillespie, 1991)

Sawmill Creek 900k Moore and Dodge, 1980

Sawmill Creek 600k Moore and Dodge, 1980

Oak Creek 1.2my Gillespie et al., 1983

82

Table 2: Summary Geochronology Key: FLVFZ, Fish Lake Valley fault zone; OVF, Owens Valley fault; PH, Poverty Hills; DSF, Deep Springs fault; WMF, White Mountain fault; OV, Owens Valley; DSV, Deep Springs Valley; BPVF, Big Pine volcanic field; DV-FCFZ, Death Valley-Furnace Creek fault zone; W-I, White-Inyo; SN, Sierra Nevada; FLV, Fish Lake Valley; B&R, Basin and Range; ECSZ, Eastern California shear zone; WL, Walker Lane Time Geologic/Tectonic Events Present FLVFZ slip rate 6-8 mm/yr; OVF slip rate 3-4 mm/yr

Late Pleisto FLVFZ slip rate ~2-3 mm/yr

0.2-0.3 Ma Red Mountain flows around PH topographic barrier

0.5-0.6 Ma basalt of E PH fault block erupted; PH uplift ongoing

0.76 Ma eruption of Bishop ash; FLVFZ lat slip rate ~11 mm/yr; DSF vert rate ~1.0 mm/yr here to present; vertical rates high on FLVFZ, DSF, OVF, WMF; rapid uplift/subsidence in OV; PH uplift ongoing; N-S drainage across DSV overcome by vertical offset on DSF (DSV opens)

1.2-1.0 Ma BPVF eruptions begin; PH uplift ongoing

1.7-1.5 Ma strain migrates from OVF thru DSF; locus of shear migrates from DV-FCFZ to OVF; slip rates decrease on WMF and DV-FCFZ; FLVFZ slip rate ~3 mm/yr, begins increasing rapidly; OVF steps westward to Crater Mtn; onset of PH transpression; onset of transtensional formation of deep graben in northern OV

3 Ma Waucoba lakebeds; period of major W-I Mtn. uplift/OV subsidence begins; Queen Valley opens (onset of strike-slip on OVF?); S OV ~1km deep;

~4.5Ma SN uplift accelerates

5 Ma central FLVFZ/northern DV-FCFZ accumulates 30 km dextral offset here to present (total 40-50 km); FLVFZ slip rate ~3 mm/yr

6-5 Ma subsidence of S OV begins; onset of formation of S OV graben (transtensional?)

6.9-4 Ma extensional faulting opens FLV; initiation of DSF

~10 Ma onset of dextral shear in W B&R; dextral motion on FLVFZ began, slip rate ~6 mm/yr; southern ECSZ, WL accumulate 60-75 km dextral offset from here until present

83

Table 3: Offset Estimates 1.7-1.5 Ma to present Northern Owens Valley fault ~2.16 km White Mountain fault 2.7-3.06 km Deep Springs fault 1.05-1.19 km Southern Owens Valley fault 5.91-6.41 km 3 Ma to 1.5-1.7 Ma Southern Owens Valley-White Mtn. fault 4.90-5.91 km Total offset 3 Ma to present 10.81-12.32 km 1.5-1.7 Ma to 5 Ma 12.44-13.79 km Total 5 Ma to present 18.34-20.2 km 1.5-1.7 Ma to 10 Ma 31.29-33.49 km Total 10 Ma to present 37.19-39.9 km

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