Abstract

The Sylvania Mountain fault system is a major left-oblique structure that extends east from the Furnace Creek–Fish Lake Valley fault in southwestern Nevada (USA). The system interacts with a series of north-northeast–striking structures that bound a rectilinear pull-apart basin, Lida Valley, and serve as part of a displacement transfer system relaying slip from the eastern California shear zone to the Walker Lane. On the basis of gravity analysis, the Lida Valley basin is internally dissected by a complex system of faults. The subsurface basin morphology differs from north to south. Major extensional faults localized displacement in the north and formed deep basins, but in the south, displacement was distributed on widely spaced structures with modest displacement. Localized extension in the north is separated from the southern domain of distributed deformation by a west-northwest oblique-slip fault. The subsurface geometry of the basin was determined from a gravity survey with measurements depth inverted in three dimensions. Geologic cross sections were constructed and their gravity signatures forward modeled for compatibility with observations. Projections of mapped faults together with structures determined from gravity modeling were combined to construct the subsurface geometry of the Lida Valley fault system and to evaluate a fault displacement budget. By conserving fault slip on the array of structures, restoration of the pre-Neogene basement to a reference datum indicates a cumulative vertical displacement of 2.3–2.5 km since the onset of basin formation. Vertical displacement estimates were used to compute the horizontal component of extension, which ranges from 1.3 to 1.4 km.

INTRODUCTION

Since the Pliocene and possibly since the mid-Miocene, displacement within the eastern California shear zone (Wallace, 1987; Dokka and Travis, 1990; Miller et al., 2001) and Walker Lane (Locke et al., 1940; Stewart, 1988; Oldow, 1992) has localized as much as 25% of the relative motion between the Pacific and North American plates (Eddington et al., 1987; McQuarrie and Wernicke, 2005; Argus and Gordon, 1991; Reheis and Sawyer, 1997; Dixon et al., 2000; Bennett et al., 2003; Oldow, 2003; Wesnousky, 2005a; Frankel et al., 2007a, 2007b, 2008, 2011; Ferranti et al., 2009; Lee et al., 2009; Ganev et al., 2010; Foy et al., 2012) as part of a major intraplate fault zone (Wesnousky, 2005b; Faulds et al., 2005). Within this tectonic belt, substantial deformation has been concentrated on the northern end of the Furnace Creek–Fish Lake Valley (FC-FLV) fault zone (Reheis and Sawyer, 1997; Reheis and Dixon, 1996; Oldow et al., 1994, 2008), that marks the eastern boundary of the northern eastern California shear zone, where displacement is relayed north-northeast to the Walker Lane (Oldow, 1992, 2003; Oldow et al., 1994, 2008). Until the mid-Pliocene the displacement transfer was accomplished on a low-angle, northwest-dipping detachment fault system underlying the Silver Peak–Lone Mountain extensional complex (Oldow, 1992; Oldow et al., 1994, 2008, 2009). Since the mid-Pliocene, younger high-angle faults have crosscut the detachment and continue to carry motion from the FC-FLV fault system to the central Walker Lane (Oldow et al., 2008; Foy et al., 2012).

In this study we assess the geometry, kinematics, and displacement budget of the Sylvania Mountain fault system, one element of this younger system of transfer faults that carries displacement from the FC-FLV fault to the central Walker Lane. In particular, we focus on the interaction between west-northwest faults with left-oblique motion and kinematically linked north-northeast extensional faults forming a rectilinear pattern of structures relaying motion to the Walker Lane by transforming the right-oblique displacement through a system of left-oblique and extensional structures. At the eastern end of the Sylvania Mountain fault system, displacement is partitioned onto a rectilinear pattern of west-northwest– and north-northeast–striking faults that bound an extensional basin underlying Lida Valley. The basin has the general geometry of an east-west–trending rhombic pull-apart basin (Burchfiel and Stewart, 1966; Crowell, 1974; Dooley and McClay, 1997; Mann et al., 1983). East-west–trending faults bounding the northern and southern basin margins are intersected at high angles by the north-northeast–striking extensional faults that form the eastern and western basin margins and dip toward the basin center. Based on a gravity data and depth modeling, we constrain the subsurface geometry of the basin to determine the degree to which known bounding faults interact with inferred subsurface structures in the basin. Using a previously established extension direction for the region (Oldow, 2003; Oldow et al., 2008; Biholar, 2011; Katopody et al., 2013), we use the vertical displacement for each inferred fault by restoration to a preextensional datum to compute the net slip on each structure and the spatial distribution of the horizontal component of slip for the basin as a whole. The slip budget is used to test the internal consistency of a geometric model for the inferred subsurface fault network, and to provide a balanced displacement budget for the basin.

REGIONAL GEOLOGIC SETTING

In the western Great Basin, contemporary displacement is transferred from the northern part of the eastern California shear zone (ECSZ; Wallace, 1987; Dokka and Travis, 1990; Ford et al., 1990; Miller et al., 2001) 75 km east to the central Walker Lane (Locke et al., 1940; Stewart, 1988; Oldow, 1992, 2003; Dickinson, 2004; Wesnousky, 2005a, 2005b) by a complex array of north-northwest, northwest, and east-northeast transcurrent and extensional faults. Transtensional deformation (Gan et al., 2000; Bennett et al., 2003; Oldow, 2003; Unruh et al., 2003; Hammond and Thatcher, 2007) in the northern part of the ECSZ) is localized on the Owens Valley–White Mountain and FC-FLV faults and is stepped east to a broad zone of northwest-striking right-oblique faults (Ekren and Byers, 1984) in the central Walker Lane. The displacement transfer system consists of a 50-km-wide belt of east-northeast–striking faults constituting the Mina deflection (Ryall and Priestley, 1975; Stewart, 1988; Oldow, 1992), which merges with right-oblique structures of the northern ECSZ and central Walker Lane (Oldow, 1992, 2003; Oldow et al., 2008; Ferranti et al., 2009) to the west and south, respectively.

Prior to ca. 3–4 Ma, before the separation of the White Mountains and the Sierra Nevada, and beginning ca. 12–10 Ma (Oldow et al., 2008), deformation in the northern ECSZ was accommodated along the northwest-striking FC-FLV fault system (Fig. 1), which extends ∼250 km from southeastern California to west-central Nevada (Stewart, 1967; Wright and Troxel, 1967; McKee, 1968; Wernicke et al., 1989) and has an aggregate right-lateral displacement of between 50 and 80 km to as much as 104 km (Stewart, 1967; Snow and Wernicke, 1989, 2000; Niemi et al., 2004; Renik and Christie-Blick, 2013). The displacement was transferred east to the belt of northwest transcurrent faults in the central Walker Lane via a low-angle detachment underlying the Silver Peak–Lone Mountain extensional complex (Oldow, 1992; Oldow et al., 1994, 2008, 2009). Activity on the detachment exhumed lower plate rocks starting ca. 12–10 Ma and displacement continued at least locally until ca. 5 Ma (Oldow et al., 1994, 2008). Activity on the detachment ceased ca. 3–4 Ma (Oldow et al., 1994, 2008) and coincided with the opening of the Owens and Chalfont Valleys in eastern California as well as with the onset of regional transtensional deformation in the western Great Basin (Henry and Perkins, 2001; Stockli et al., 2000, 2003; Oldow et al., 2008).

Ancient and contemporary displacement on the northern part of the FC-FLV fault system decreases from south to north, both in total magnitude and rate. In the northern parts of the fault system, displacement in the northern Death Valley and Sylvania Mountains (Fig. 1) is estimated as 50 km on the basis of offset segments of Jurassic plutons (McKee, 1968). Farther north in Fish Lake Valley, offset of a Paleozoic thrust is reduced to ∼20 km (Renik and Christie-Blick, 2013). Similarly, late Pleistocene to recent deformation rates, determined by offset of radiometrically dated geomorphic surfaces (Reheis and Sawyer, 1997; Frankel et al., 2007a, 2007b; Ganev et al., 2010) decrease from 6 mm/yr for the Furnace Creek fault in the Sylvania Mountains to ∼2 mm/yr for the central and northern Fish Lake Valley fault farther north. The discrepancy in cumulative displacement was attributed to the transfer of motion east to the central Walker Lane via the Silver Peak–Lone Mountain extensional complex (Oldow et al., 1994, 2008, 2009). The south to north reduction in Quaternary slip rate, however cannot be accommodated by the Silver Peak–Lone Mountain detachment, which ceased activity by 3–4 Ma, and requires displacement transfer on younger structures (Frankel et al., 2007a, 2007b; Foy et al., 2012).

Expanding on the work of Foy et al. (2012), our ongoing mapping and structural analysis in the northern part of the southern Walker Lane (Oldow and Geissman, 2013) has identified a complex array of active high-angle faults that link the northern FC-FLV fault system with transcurrent faults of the central Walker Lane. The high-angle faults crosscut and postdate low-angle detachment structures in the region and transfer displacement east from the FC-FLV and north to the central Walker Lane (Fig. 1). The active faults form a belt of structures ∼20 km wide and 70–100 km long defined to the north and south by west-northwest–striking anastomosing fault systems with as much as 30 km of aggregate left-lateral displacement (Oldow and Geissman, 2013; Katopody et al., 2013). The northern and southern fault zones are the Palmetto Mountain and Sylvania Mountain fault systems, respectively, and are linked by a diffuse zone of faults that underlies the intervening region. From west to east, the bounding fault zones change strike from west-northwest to east-west and at a distance of ∼30–35 km from their western intersection with the FC-FLV fault merge with north-northeast–trending extensional faults that extend north to the eastern part of the central Walker Lane (Fig. 1). The extensional faults bound a series of basins and ranges as much as 40 km wide and successively accommodate left-oblique motion from the bounding faults (Oldow and Geissman, 2013; Katopody et al., 2013; Dunn et al., 2013; Nix et al., 2013).

The Sylvania Mountain fault system marks the southern margin of the active displacement transfer system, as a belt of anastomosing west-northwest– to east-west–striking faults. The fault system extends east for nearly 70 km, to the eastern physiographic boundary of the southern Walker Lane (Stewart, 1988) and ∼25–30 km east of the FC-FLV fault zone broadens from 6 to 8 km where fault strands dissect and separate Slate Ridge from Lida Valley to the north. North-northeast–striking faults that bound the eastern and western basin edges beneath Lida Valley merge with and link the west-northwest–trending segments defining the northern and southern basin margins.

The Sylvania Mountain fault system cuts a low-angle detachment that underlies both Slate Ridge and the Sylvania Mountains. The detachment records large-magnitude extension involving pre-Cenozoic and Miocene rocks (Oldow and Geissman, 2013) and was active following deposition of ca. 11.3 Ma ignimbrites from the Timber Mountain volcanic center. The detachment is locally overlain unconformably by welded tuff dated as 7.6 Ma (Noble et al., 1964; Weiss et al., 1993) and older interbedded volcanic and volcaniclastic rocks, dated as 8.5 Ma (Verdel and Stockli, 2011), deposited in a syntectonic half-graben. The detachment is offset by strands of the Sylvania Mountain fault system. By correlating offset segments of the low-angle fault, a minimum left-lateral component of displacement of 15 km is established for the younger high-angle fault system (Oldow and Geissman, 2013).

Displacement on the Sylvania Mountain fault system is locally transferred, at least in part, to high-angle structures in and around Lida Valley. Displacement on the high-angle structures initiated by the late Miocene. Most of the structures have poorly preserved fault scarps, but some scarps suggest Pleistocene activity. No clear evidence of Holocene motion is observed on any faults.

STRUCTURE AND STRATIGRAPHY OF THE LIDA VALLEY BASIN

Lida Valley (Fig. 2) is a topographic depression located at the eastern extent of the Sylvania Mountain fault zone, where west-northwest fault strands merge with north-northeast faults that bound and cut across the valley. Lida Valley is flanked by Magruder Mountain and Slate Ridge to the west and south, respectively, by Mount Jackson Ridge to the north, and to the east by a low high named here the Cottontail Ranch Hills. The rectilinear basin extends east-west for ∼25 km and north-south for 10 km. The valley floor is covered by Quaternary alluvium and decreases west to east from a maximum elevation of 1900 m to 1500 m at the lowest point of the depression (Fig. 2). The elevation difference between the valley surface and adjacent hills and mountains ranges from 100 to 950 m.

The most prominent topographic highs are Magruder Mountain at 2750 m, which forms an abrupt escarpment that decreases south to north from 2750 to 2200 m along the western boundary of Lida Valley, and Mount Dunfee to the southeast. Mount Dunfee at 2100 m is the highest point on the east-west–trending Slate Ridge. Slate Ridge has a subdued topographic relief of 1600–2100 m and extends west 30 km to the eastern flank of the Sylvania Mountains, which have elevations of 1900–2100 m. The east-west–trending Mount Jackson Ridge forms the northern flank of the valley and has elevations of 1700–2000 m. The eastern basin margin consists of a gradual topographic transition from the valley to the Cottontail Ranch Hills with elevations of ∼1650 m.

Fault Geometry

The basin is bounded by a rectilinear array of extensional and left-oblique faults with north-northeast and west-northwest strikes, respectively (Fig. 2). The transition between north-northeast extensional faults and west-northwest left-oblique faults is well exposed in several locales where the faults merge through curved arrays of structures with intermediate orientations. The Lida Valley basin is separated from the surrounding highlands by these faults with west-northwest structures defining the north and south boundaries and north-northeast structures defining the west and east boundaries (Fig. 2).

West-northwest–trending strands of the Sylvania Mountain fault system enter Lida Valley at the southwest. The fault zone is ∼6 km wide and is composed of three major fault strands, locally connected by anastomosing structures. The southern and central strands are mapped east into and adjacent to Slate Ridge, where the southern strand can be traced east to the south side of Mount Dunfee. The central strand projects into the southwestern part of Lida Valley and is locally identified by scarps in older alluvial deposits. It is mostly covered by younger alluvial deposits and is approximately located by the abrupt end of pre-Cenozoic outcrops. The northern strand of the Sylvania Mountain fault emerges from the northern side of the topographic divide between the Sylvania Mountains and Magruder Mountain, and bifurcates with one trace continuing east where it is lost beneath younger alluvium, while the second segment curves north to merge with the East Magruder Mountain fault (Reheis and Noller, 1991; Piety, 1996).

The western margin of the basin is marked by north-northeast faults that track along the eastern flank of Magruder Mountain (Fig. 2). The East Magruder Mountain fault is major structure and is mapped as a series of discontinuous segments, locally covered by alluvial deposits, particularly at the mouths of major canyons, and consists of two parallel strands recording down-to-the-east displacement into the valley. The fault system emanates from the northernmost strand of the Sylvania Mount fault system, exposed ∼3 km north of the headwaters of Tule Canyon in the highlands along the southwestern part of Lida Valley (Fig. 2). The kinematic link between the East Magruder Mountain fault and the Sylvania Mountain fault system is well expressed as a series of discontinuous fault segments that curve through ∼80°–90° from west-northwest to north-northeast. The north-northeast range front fault is tracked nearly 8 km to the northwest extent of Lida Valley, where the valley is flanked by low hills to the east that expose Paleozoic carbonate and clastic rocks constituting a triangular topographic protrusion into the valley. The Paleozoic rocks in this basement exposure are overlain by basalt and are bounded on the west and east by west-dipping and east-dipping high-angle faults, respectively.

The northern end of the East Magruder Mountain fault merges with a complex array of east-west– to east-northeast–striking faults that mark the northern margin of Lida Valley. The east-west fault system is composed of at least two strands of subparallel structures that are discontinuously mapped to the east along the southern flank of Mount Jackson Ridge. The faults show down-to-the-south displacement and extend discontinuously east to the low topographic divide that forms the northeastern margin of Lida Valley. The geometry of the faults along the northern boundary of the valley is complex. The northern zone of east-west faults is an eastern extension of northeast-striking faults that bound Lida Wash, the narrow northeast-trending topographic trough that separates Magruder Mountain from the Palmetto Mountains to the north (Fig. 2). The northern zone of faults in the northwestern part of Lida Valley have a dogleg morphology that offsets the axis of western Mount Jackson Ridge to the north. To the east, the faults track the southern margin of the ridge, and are linked to north-northeast–trending high-angle faults bounding ridges of pre-Cenozoic layered rocks overlain by Cenozoic volcanic rocks. The east-west system of faults is segmented by two northeast-striking faults, both with down-to-the-east displacement. The westernmost northeast-striking fault is an extension of the structure that bounds the eastern margin of Paleozoic rocks in northwestern Lida Valley. The eastern northeast-striking structure is mapped across Mount Jackson Ridge east of Mount Jackson (Fig. 2) and continues to the north-northeast along a ridge of Paleozoic rocks overlain and intruded by rhyolite.

The eastern boundary of Lida Valley is fairly simple and consists of the north-northeast–trending Jackson Wash fault that emerges from the central strand of the Sylvania Mountain fault system northeast of Mount Dunfee. The fault marks the western extent of bedrock exposures along the eastern side of the basin, and with the exception of a few short segments where alluvium is juxtaposed against Paleozoic clastic rocks, is obscured by younger alluvial deposits. Farther east, in the Cottontail Ranch Hills, a subparallel down-to-the-west fault also emanates from the eastern continuation of the central Sylvania Mountain fault system along the north side of Mount Dunfee and eastern Slate Ridge.

Stratigraphy

The mountains and ridges flanking Lida Valley expose Proterozoic and Paleozoic carbonate and clastic rocks, Jurassic granitic plutons, and late Cenozoic volcanic and sedimentary rocks (Albers and Stewart, 1972). The pre-Cenozoic carbonate and clastic rocks are locally intruded by Jurassic granitoids and overlain with angular unconformity by late Miocene silicic and mafic volcanic rocks and interbedded volcaniclastic sedimentary rocks (Albers and Stewart, 1972). The distribution of the many lithologic constituents varies around the basin, with Paleozoic carbonate and clastic rocks best exposed in Magruder Mountain to the west and around Mount Dunfee to the southeast (Fig. 2). Mount Dunfee and the Cottontail Ranch Hills expose Proterozoic and Paleozoic carbonate and clastic rocks overlain by Cenozoic volcanic and sedimentary rocks. Mount Jackson Ridge exposes a sequence of interstratified Cenozoic volcanic and sedimentary rocks overlying Paleozoic carbonate and clastic rocks with angular unconformity. With the exception of Mount Dunfee in the east, where Paleozoic and Proterozoic rocks are exposed, Slate Ridge is composed of Jurassic granitic plutons containing screens and roof pendants of Proterozoic clastic and carbonate rocks with scarce exposures of Cenozoic volcanic rocks.

Cenozoic rocks are exposed on the north, south, and east flanks of Lida Valley but are not preserved on Magruder Mountain. The thickness of the volcanic and sedimentary rocks varies from 120 to 150 m, and is characterized by significant lateral differences in stratigraphy, possibly reflecting variations in volcanic source, probable topographic control of deposition, differential preservation, and the effects of late Cenozoic deformation. The Cenozoic stratigraphy is divided into two stratigraphic sequences that overlie Slate Ridge, the Cottontail Ranch Hills, and Mount Jackson Ridge. The lower sequence consists of interbedded air-fall and ash-flow tuff and a locally exposed unit of andesite flows and lahar found at the base of Mount Jackson (Fig. 2). The upper sequence is composed of interbedded volcanic tuff, basalt, and volcaniclastic rocks that are intruded by 2.9 Ma rhyolite plugs and domes (Mckee et al., 1989). The upper sequence overlies rocks of the lower sequence and pre-Cenozoic layered and plutonic rocks.

The lower sequence ranges in thickness from 50 to 150 m and overlies, with angular unconformity, Paleozoic rocks in Mount Jackson Ridge and in the eastern parts of Slate Ridge. In much of the Cottontail Ranch Hills and Slate Ridge, lower sequence rocks are separated from underlying Proterozoic strata and Jurassic plutonic rocks by a low-angle detachment (Oldow and Geissman, 2013). The basal unit of the lower sequence is exposed at Mount Jackson and consists of an undated succession of tuff and andesite, with ∼50 m of andesite lying between upper and lower tuffs that are both ∼100–150 m thick. The age of these volcanic rocks, which are thought to be the lowest part of the sequence, is based on correlation with lithologically similar rocks elsewhere. The lower tuff may correlate with a 21 Ma tuff (Robinson et al., 1968) overlain by andesite in the western Silver Peak Range and the upper tuff may correlate with the basal unit of the lower sequence exposed in the Cottontail Ranch Hills and Mount Dunfee region. In these areas to the east and south, the lower sequence consists of 50–80 m of interleaved tuff with a basal unit dated as 16.5 passing upward to tuffs, dated as 13.9 Ma (Noble et al., 1993). Tuff thicknesses vary due to multiple internal unconformities within the succession (Noble et al., 1993) and virtually all units are overlain by members of the Timber Mountain tuff sequence dated as between 11.3 and 11.6 Ma (Sawyer et al., 1990) that, although separated by an angular unconformity (Noble et al., 1993), is included in this stratigraphic sequence.

The upper sequence ranges in thickness from 60 to 110 m and varies substantially in lithologic constituents and thickness across the study area. Compositionally the unit contains sedimentary rocks, ash-flow tuff, and basalt, but the stratigraphic succession varies. At Mount Jackson Ridge, the unit is composed of a basal unit as much as 40 m thick of ash-flow tuff (Stonewall Flat) dated as 7.6 Ma (Sawyer et al., 1990) that is locally overlain by an undated basalt as thick as 40 m. Both successions are overlain unconformably by 2.9 Ma rhyolite flows (McKee et al., 1989) as thick as 80 m. The rhyolite flows are the extrusive equivalent of a rhyolite dome and plug field found on Mount Jackson Ridge and the ranges to the northeast, where the intrusive rocks and extrusive equivalents intrude or overlie Paleozoic strata and are localized along north-northeast–striking high-angle faults. Elsewhere, in the Cottontail Ranch Hills and Slate Ridge, the basal tuff overlies volcanic rocks of the lower sequence and Proterozoic and Paleozoic strata and granitoids of the basement. In a few locations, interbedded volcanic and volcaniclastic sedimentary rocks locally dated as 8.5 Ma (Verdel and Stockli, 2011), and as thick as ∼100 m, are below the basal tuff. In the western reaches of Slate Ridge, undated basalt overlies Jurassic granitoid with nonconformity. Elsewhere, north of Magruder Mountain and at Black Mountain, on the southern basin edge, the basalt unit is as much as 70 m thick.

Subsurface Basin Morphology

Observed Gravity

To describe the subsurface geometry of Lida Valley, gravity data were collected and a residual complete Bouguer anomaly (RCBA) was generated. Prior to this study, data coverage for Lida Valley and the surrounding region, provided by the Pan American Center for Earth Studies (PACES), consisted of 14 gravity stations within the basin and 60 measurements in the adjacent mountain ranges and intervening valleys. To improve spatial coverage we collected ∼500 gravity measurements at 300 m spacing along 7 transects crossing Lida Valley (Fig. 3). Access was limited by dense sagebrush and, for the most part, the lines followed primary and secondary roads and in a few locations were acquired by tracking cross country. Four transects consist of lines originating and terminating in bedrock (lines 1 through 4), and two are controlled by bedrock at one end (lines 6 through 7). Several other lines, including lines 5 and 8, link the primary transects and in some cases are discontinuous and extend into the basin as far as access allowed. To ensure internal data consistency, at all line intersections a common station was reoccupied. Four transects extending between bedrock exposures consist of two north-northeast lines (lines 3 and 4) following the western and eastern margins of the basin, respectively, a west-northwest–trending line (line 2) along the southern margin of the basin, and a northwest-trending line (line 1) passing diagonally across the basin. An east-west–trending transect (line 6), tracking along the northern margin of the basin, extends from bedrock exposures in the west 19 km east into the western part of Stonewall Flat. This line was linked to basement exposures along Mount Jackson Ridge to the north by the northern extent of line 4 in the west and in the east by line 8. Two curved lines (lines 5 and 7) from bedrock exposures in the southwest corner of the basin extend across the central portion of the basin and intersect the east-west–trending line 6 to the north.

The data were collected using two Scintrex CG-5 gravimeters with station locations determined using dual-frequency Leica GS10 global navigation satellite system (GNSS) receivers. Each gravity station value consists of the average of three 60 s measurements, each composed of the mean of 1 s determinations of the relative gravity. All 500 station values were referenced to a common base station that was measured by both gravimeters at the beginning and the end of each day. A secondary base station, located at the GNSS base station, was also measured by both gravimeters daily. The gravity base station was referenced to an absolute gravity station, Las Vegas K 169, located at the southwest corner near the front entrance of the United States Post Office in Las Vegas, Nevada. This post office was decommissioned in 2012 and is now a museum, but access to the absolute gravity monument is still possible. The GNSS position data were collected in the real time kinematic (RTK) mode and were post-processed using Leica GeoOffice. The GNSS base station was located on a bedrock outcrop 1.6 km north-northwest of Gold Point, Nevada, along the southern margin of Lida Valley. Relative position uncertainties between the base station and rover ranged from 0.016 to 0.01 m. All GNSS positions of the survey were georeferenced in an Earth centered Earth fixed (ECEF) frame provided by the Continuous Observation Reference System (CORS) daily for 11 days using the Online Positioning User Service (OPUS) supported by the National Geodetic Survey (NGS). The base station was located at an ECEF reference position of lat 37°21′27.38084″, long 117°22′14.99551″ with a maximum uncertainty of 1.3 cm (±0.0017), used to transform all relative positions to a common ECEF frame.

A complete residual Bouguer anomaly was computed using measurements from our gravity survey combined with ∼65,000 values from the Pan American Center for Earth and Environmental Studies (PACES) (http://research.utep.edu/Default.aspx?tabid=37229, 2005) gravity database (to a radius of 166.7 km) from Lida Valley. In conformity with the new standards set by the U.S. Geological Survey (USGS) (Hildebrand et al., 2002) and the Standards/Format Working Group of the North American Gravity Database Committee (Hinze, 2003), the Bouguer anomaly computations are based on ellipsoidal heights (Holom and Oldow, 2007). All gravity data for our survey are referenced to ellipsoidal heights directly acquired from GNSS positioning. In contrast, PACES data are referenced to orthometric heights and required transformation to the ellipsoid using the NGS National Oceanic and Atmospheric Administration transformation program Geoid09 (http://www.ngs.noaa.gov/GEOID/GEOID09/). The consistency between our measurements and the PACES values within the basin was determined for the 14 stations within Lida Valley. Four PACES stations were rejected on the basis of inconsistency with a regional complete Bouguer anomaly computed solely with PACES data. For 10 stations, located within 50–350 m of our survey stations in the low-relief parts of the valley, the computed absolute gravity values are within 0.1–0.3 mGal and were combined with the results of our survey. Our gravity data were terrain corrected (Cogbill, 1990) and combined with the height- and terrain-corrected PACES data in an Excel spreadsheet (Holom and Oldow, 2007) to produce a regional complete Bouguer anomaly using a reduction density of 2.67 g/cm3.

An RCBA was computed for the basin. The irregularly spaced complete Bouguer anomaly values were gridded using a minimum curvature cubic spline algorithm with no tension that honored all data points. We used the Mickus et al. (1991) technique, consisting of the difference between the interpolation including and excluding the data within Lida Valley to produce the residual gravity signature.

Lida Valley is underlain by distinct west-northwest–trending RCBA lows in the north and the east and by low-amplitude highs and lows in the southwest and southeast (Fig. 3). Given the distribution of the survey lines, several of the prominent features are well constrained but others are not. In the southwestern part of the basin, from east to west a –3 mGal low is bounded by a gravity high of +1 mGal that extends across the central axis of this part of the basin. Farther west, a west-northwest–trending low-amplitude high separating two well-constrained lows of –1 to –2 mGal is probably an artifact of grid interpolation. Toward the northeast from the southwest part of the basin, the gravity ranges from –2 to –4 mGal along a west-northwest trend before entering the west-northwest gravity low characteristic of the northern and eastern segments of the valley. The north-northwest–trending gravity low has a maximum value of –9.3 mGal and is bounded on the west by a steep gradient of –8 mGal to +1 mGal over 2 km. The southern extension of this western gradient is poorly constrained by the data distribution, as is the morphology of the southwestern margin of the low. The gravity signature of the eastern boundary of the west-northwest–trending low decreases gradually in magnitude from –8 to 0 mGal over 5 km, with a steeper portion of the gradient of –8 to –3 mGal over a distance of 3 km. The eastern boundary of the gravity low is controlled by limited data, but coincides with the edge of the basin, which is confirmed by exposures of bedrock. Internally, the gravity low is divided by a modest, north-northeast–trending relative high (between –7 and –8 mGal) forming a saddle morphology. The northern margin of the western part of the low is well constrained as –9 to –3 mGal over a distance of 2 km. The shallow gradients for the northern and southern boundaries of the eastern part of the low, however, are poorly controlled by the data distribution and possibly reflect an interpolation artifact.

Depth Inversion

As a first step toward modeling the basin, the RCBA was inverted for depth to basement in three dimensions using the Geosoft Oasis Montaj GM SYS-3D modeling software (http://www.geosoft.com/products/gm-sys/gm-sys-3d-modelling), which uses the Parker-Oldenburg algorithm for gravity inversion (Oldenburg, 1974). In that there are no previous density-depth studies from Lida Valley, we used upper and lower bounds for constant density models set at 2.2 g/cm3 and 2.4 g/cm3, respectively. The lower bound for density of 2.2 g/cm3 was derived from boreholes in alluvial and lacustrine deposits ∼1.4 km thick in Hot Creek Valley, located 165 km north-northeast of Lida Valley (Healey, 1970). To account for the possibility of water-saturated basin fill and presuming a sediment porosity of 20%, an upper bound of 2.4 g/cm3 was employed. The inversion process requires the construction of three spatially coincident grids constrained to the boundary between bedrock and basin fill. In our study, three grids with a 300 m cell size were used and consisted of (1) topography interpolated from the elevation control for the gravity stations, (2) an assigned density distribution, and (3) interpolated RCBA values. The topographic grid was generated by interpolating the ellipsoidal heights measured for the gravity stations and was compared to the USGS National Elevation Dataset (NED) 10 m resolution digital elevation model (DEM) to test for blunders in our surveyed elevations. We transformed the USGS NED 10 m DEM from orthometric to ellipsoidal heights to be consistent with our DEM. The USGS NED 10 m DEM was resampled at 300 m cell size and differenced with our DEM. The residual results indicated that there were no discrepancies at a resolution of 10 m. The interpolated gravity data were upward continued to a horizontal datum (Cordell, 1976) that corresponds to the maximum elevation 300 m above the lowest elevation within the basin. This resulted in a maximum of loss of RCBA gravity signal of 2 mGal.

The depth inversion models (Figs. 4A, 4B) bear a close similarity to the RCBA in the northern and eastern portion of the basin but show significant differences in the southwestern and southeastern quadrants. In the north, both models clearly delineate the two subbasins previously inferred from the RCBA and recognize the modest north-northeast–trending high that creates a saddle morphology. In the first inversion, using a constant density for basin fill of 2.2 g/cm3, maximum depths for the eastern and western subbasins of ∼350 m and ∼365 m, respectively, are computed, and are divided by a subtle high with an estimated subsurface depth of ∼320 m. The model preserves the orientation of the northeast-trending gravity low, represented by a –3 mGal contour in the RCBA, along the south-central basin margin; however, the alternating low-magnitude highs and lows observed in the RCBA in southwestern corner of the basin and along the southern margin are absent due to smoothing by inversion surface. The second inversion model, using a constant density of 2.4 g/cm3 for basin fill, yields maximum depths for the eastern and western subbasins at ∼610 m and ∼625 m, respectively, and the north-northeast–trending high separating the basins has a computed depth of ∼565 m. Unlike the first inversion, the subsurface morphology of the two subbasins is accentuated by prominent gradients in the interpolated surface. Some of the smaller magnitude alternating highs and lows present in the RCBA in the southwestern portion of the basin and along the southern margin remain absent in the smoothed surface of the second inversion.

Subsurface Architecture

The location of major subsurface faults with little or no surface expression within and bounding the basin can be inferred from gradients in the RCBA and depth derived from the inversion models. The basin is divided into three domains (Fig. 4C) that display different gravity signatures and significant variations in depth that are bounded by corresponding gradients in the RCBA and depth inversions. The western domain underlies the entire western third of the basin and is characterized by subdued gravity gradients and relatively shallow depth to basement. The eastern two-thirds of the basin are underlain by the northern and southern domains, which are divided by a west-northwest boundary. The northern domain is characterized by two circular negative gravity anomalies that correspond to two deep depressions in the basin floor, in sharp contrast to the southern domain, which is characterized by a pattern of alternating gravity highs and lows with modest amplitudes.

In the northern domain, the two equant lows in gravity and corresponding basin deeps derived from depth inversion models are bounded to the west by a steep east-facing gradient that trends north-northeast, on the east by a modest west-facing gradient, and on the north and south by east-west– to west-northwest–trending gradients. The western boundary is interpreted as a major east-facing fault system, as is the eastern flank of the intervening north-northeast–trending high that divides the depression. The eastern flanks of both depressions are modeled as west-dipping ramps, with minor west-facing extensional faults in the eastern part of the domain. The northern margin of the depressions is defined by a prominent gradient of both the RCBA and depth over a distance of ∼3.5 km to exposures of bedrock in the Mount Jackson Ridge and marks the location of the exposed and inferred Mount Jackson Ridge fault system. The boundary between southern and northern domains is totally within Lida Valley and has no surface expression. The boundary trends west-northwest, is defined by a north-facing gradient in gravity and depth, and is interpreted as a fault system with down-to-the-north displacement.

The southern domain displays a gravity signature of alternating modest highs and lows that are well defined in the RCBA (Fig. 4C) but more subtle in the depth inversions (Figs. 4A, 4B) and is flanked by a relatively steep north-facing gradient that follows the boundary between Slate Ridge and Lida Valley. The gradients are interpreted as a series of east- and west-dipping low-magnitude faults striking north-northwest that produce horst and graben structures. The gradient along the southern boundary is ∼1.7 km wide and coincides with the projection of mapped segments, east and west, of the central Sylvania fault system (Fig. 2).

The western domain differs from the domains farther east by having both modest negative and positive gravity anomalies and a shallow estimated depth to basement. Unlike the areas in the main part of Lida Valley, several active faults are mapped and have orientations consistent with the Sylvania fault system in the south and with north to north-northeast trends that reflect the fault discontinuously exposed along the east flank of Magruder Mountain, which forms the western boundary of the domain. Depth inversion estimates of a shallow basement for the domain are supported by three outcrops of Pliocene basalt in the southern part of the domain (Fig. 2) and by the positive gravity anomaly that forms the north-south–trending eastern boundary of the domain. The gravity high is carried south into the subsurface from exposures of Paleozoic rocks (Fig. 3) with densities higher than the 2.67 g/cm3 used in gravity reduction. The Paleozoic rocks are highly mineralized and the southern projection of the gravity high is thought to correspond to a southern continuation of the mineralized rocks into the shallow subsurface. The exposures of Paleozoic rocks are bounded by a west-facing fault to the west and by the major gradient marking the western margin of the northern domain to the east. The exposures and subsurface positive gravity anomaly are modeled as two extensional faults that converge to the south.

Two-Dimensional Forward Models

The subsurface geometry of the Lida Valley basin was further constrained by a series of two-dimensional (2D) forward models (sections 1–8 in Fig. 5). The 2D forward models provide critical constraints on the location and facing of inferred subsurface faults, particularly for areas (such as the southwestern quadrant of the valley) where structural complexity is insufficiently imaged by the data coverage.

The 2D forward models are based on geologic cross sections with basin depth constrained by the 3D inversions as initial input. Where misfits were recognized between the predicted gravity signature of the geologic models and observed gravity, the geologic sections were modified and the gravity signature recomputed iteratively until solutions with residuals of ∼0.1 mGal were obtained. Although containing uncertainties, these models incorporate geologic relations gleaned from our understanding of the structure and stratigraphy of the study region and help improve basin depth assignments over those determined from the 3D inversions.

There are several limitations imposed on the gravity modeling. The 2D assumption is typically violated by the basin geometry and conspires to minimize depth determinations by the contributions from off-axis bodies. To mitigate these effects, we used 2.75D modeling in areas where the 3D approximation was clearly untenable. Direct determination of subsurface fault dips cannot be resolved with the acquisition spacing of 300 m. We assigned dips of 60° and located the position of the fault as the inflection of observed gravity gradients in the sections. The gravity gradients inferred to represent major faults were identified by significant gradients in the interpolated RCBA and, where they crossed lines at an oblique angle, the apparent dip of the fault was used in the forward models. Shallow gravity gradients were modeled using modest undulations or ramps in the basement structure as a means of limiting the incorporation of faults in the structural models. Constant densities were used for geologic units modeled (Table 1) and were assigned on the basis of average densities defined by Maxant (1980). In a few locations, regions of elevated density were invoked for basement rocks (>2.67 g/cm3) where positive anomalies were observed in the vicinity of known or inferred zones of mineralization.

Section 1. Section 1 (Fig. 5) trends northwest-southeast along the Lida Road and extends from the mouth of Lida Canyon in the northwest across Lida Valley south to the western flank of Mount Dunfee (Fig. 3), obliquely crossing both the northern and southern domains. There is a pronounced gravity gradient in the RCBA decreasing from +6 mGal to –9 mGal west to east (Fig. 3) at the northwest margin of the basin. This steep gradient is modeled as 3 down-to-the-east extensional faults with vertical displacements of 230 m, 440 m, and 200 m, respectively. These structures are modeled with apparent dips consistent with the model line and the assigned 60° true dip for each structure. At the boundary between the north and south domains, section 1 crosses a west-northwest–trending gradient (Fig. 3), which is modeled as a north-facing with of ∼100 m of throw. Within the southern domain, section 1 crosses a north-northeast–trending gradient that increases northwest to southeast from –3 mGal to +1 mGal (Fig. 3). This gradient is modeled as 3 down-to-the-west extensional faults with vertical offsets, from west to east, of 50 m, 100 m, and 110 m, respectively. Section 1 crosses the prominent gradient along the southern flank of the basin and is modeled as a north-facing structure with 230 m of throw. The section continues south of the town of Gold Point along Lida Road (Fig. 3) and crosses a shallow half-graben with an estimated maximum depth of ∼300 m. Exposures of Proterozoic metasedimentary rocks are modeled at the same density as bedrock, 2.67 g/cm3, in the 2D profiles. The basalt is not included within the basin model south of the central strand of the Sylvania fault system because there is no evidence of preservation in this part of Slate Ridge.

Section 2. Section 2 (Fig. 5) trends east-west from the eastern edge of Magruder Mountain to the town of Gold Point and crosses the western and southern domains. The western end of the profile extends west beyond the subsurface expression of the north-northeast–trending East Magruder Mountain fault zone (Fig. 2), which in our model is shown as a single structure having a bedrock offset of 170 m. A shallow gradient, decreasing from 0 mGal to –2 mGal to the east (Fig. 3), is modeled in section 2 as an east-dipping ramp extending from the East Magruder Mountain fault to a down-to-the-east extensional fault bounding the western side of a shallow graben with an approximate depth of 150 m. The west bounding fault dips 60° and reflects a throw of 100 m. Farther east, the gravity gradient increases from –2 mGal to +1 mGal (Fig. 3) and is modeled as a west-dipping fault with a displacement of 90 m. In the footwall of this structure, Pleistocene age basalt is exposed and corresponds with the gravity high. We invoke mineralized Paleozoic rocks for the basement to accommodate the positive gravity anomaly (Fig. 3), which is consistent with a southern projection of mineralized Paleozoic rocks exposed to the north. To the east section 2 enters the southern domain, which is characterized by a series of gravity highs and lows with values ranging from +1 mGal to –4 mGal (Fig. 3). These short-wavelength changes in gravity are modeled as alternating horst and graben structures bounded by east-dipping and west-dipping extensional faults with apparent dips of 55°. The faults have displacements ranging from 50 m to 210 m. Three down-to-the-west normal faults are modeled in the eastern half of section 2, just northwest of the town of Gold Point. These three structures are also encountered in section 1 to the north, where they have the same facing and offsets. At the eastern end of section 2, the line crosses a prominent east-west–trending gradient modeled as a north-facing structure with a vertical offset of ∼250 m.

Section 3. Section 3 (Fig. 5) trends north-south and extends from the southern basin-bedrock transition at Slate Ridge to the highlands at the northwestern corner of Lida Valley. This line extends along the axis of the western domain (Fig. 4C). The southern end of the profile is anchored in Sylvania pluton (Fig. 2) granitic rocks, which are in the lower plate of a detachment fault exposed in the headwaters of northern Tule Canyon. From south to north, the gravity gradient decreases from 0 mGal to –2 mGal across a north-facing fault with 140 m of throw, but abruptly increases to +1 mGal (Fig. 3) where basalt is exposed. The basalt is bound by a south-facing fault with an offset of ∼30 m. Continuing north along section 3, the gravity values decrease to a 0 mGal contour which extends for only 500 m and then abruptly increases to +6 mGal (Fig. 3) where the profile flanks the western edge of a Paleozoic outcrop that has undergone extensive mineralization. The northern end of section 3 is modeled as a shallow graben bounded by a north-facing fault with an offset of ∼100 m and a south-dipping fault with a throw of 170 m, on the south and north, respectively. Farther north, where the section flanks the triangular protrusion of Paleozoic bedrock, the gravity anomaly increases to +6 mGal (Fig. 3) and is associated with an extensive zone of mineralization in the highlands bounding the northwest corner of the basin.

Section 4. Section 4 (Fig. 5) begins in the town of Gold Point and extends north following Nevada State Highway 774, crossing the eastern parts of the southern and northern domains before terminating in Mount Jackson Ridge (Fig. 3). The east-west–trending gradient, decreasing to the north from +1 mGal to –3 mGal (Fig. 3), is modeled as a north-dipping structure with an offset of 280 m. This fault marks the boundary between Lida Valley and Slate Ridge and is the same structure modeled in sections 1 and 2. Farther north, section 4 crosses a prominent west-northwest–trending gradient decreasing from –2 mGal to –8 mGal that marks the boundary between the southern and northern domains (Fig. 4C). This gradient is modeled as a north-facing structure with an offset of 150 m. To the north, a northeast-trending gradient is modeled as a major west-dipping extensional fault with an offset of 240 m. Continuing northeast, a gradual gradient of –8 mGal to –2 mGal over ∼4.5 km corresponds to the eastern flank of Lida Valley and is modeled as a shallowly west-dipping ramp (Fig. 3). The northern extent of section 4 continues to Mount Jackson Ridge, where it crosses an east-west–trending gravity gradient increasing south to north from –2 mGal to +3 mGal over a distance of ∼2 km (Fig. 3). This gradient is modeled as two east-west–trending faults that define the northern basin boundary.

Section 5. Section 5 (Fig. 5) is not tethered to bedrock at either end, and trends north-south along the subtle gravity high separating the two subbasins in the northern domain. The line intersects sections 2 and 6 at the southern and northern ends, respectively (Fig. 3). It crosses a prominent east-west–trending gradient modeled as a north-facing structure with an offset of ∼120 m that corresponds in location to sections 1 and 4. Continuing north, a second east-west–trending structure is modeled with an offset of 100 m. This structure is not required by the gravity signature, but is incorporated to remain consistent with the projection of the curved structure modeled in section 1. Farther north, the section obliquely crosses a modest north-south–trending gradient (Fig. 3) modeled as graben bounded by low-angle north- and south-dipping structures with vertical offsets of 280 and 310 m.

Section 6. Section 6 (Fig. 5) trends east-west across the northern boundary of the northern domain following Nevada State Highway 266 from the mouth of Lida Canyon, crossing the entire extent of the basin and, continuing beyond Lida Valley into Stonewall Flat (Fig. 3). The pronounced gradient at the northwest corner of the basin (Fig. 3) is modeled as two major down-to-the-east extensional faults with vertical offsets of 230 m and 440 m. A third east-dipping fault is required to fit the observed gravity gradient and has a vertical displacement of 570 m. Continuing east, section 6 crosses the deepest portion of the western subbasin and trends east across a north-south gradient that increases from –9 mGal to –3 mGal over ∼2.5 km (Fig. 3). This gradient is accommodated by a west-dipping ramp bounded on the eastern side by a major east-dipping normal fault with an aggregate displacement of ∼350 m. Farther east, section 6 crosses a series of very shallow gravity highs and lows (Fig. 3) modeled as highly oblique intersections of east- and west-dipping faults with vertical offsets from 170 m to 210 m.

Section 7. Section 7 (Fig. 5) trends northeast from the southern basin fill–bedrock transition just north of Tule Canyon, obliquely crossing the southern part of the western domain, the northwestern segment of the southern domain, and passing to the center of the basin, where it intersects section 1 (Fig. 3). The southern extent of the profile overlies the Sylvania pluton (Fig. 2), which extends north in the subsurface in a manner similar to that depicted in section 3. The profile obliquely crosses a southwest–trending gradient that decreases from 0 mGal to –3 mGal (Fig. 3) and to the northeast crosses 3 low-magnitude gravity highs and lows. The gradients are modeled as a series of east- and west-dipping extensional structures recording vertical offsets ranging from 180 to 140 m. The northern end of the profile crosses the prominent west-northwest gradient (Fig. 3) visible in sections 1, 4, and 5, and is modeled as a north-facing extensional structure with a vertical offset of ∼100 m.

Section 8. Section 8 (Fig. 5) trends south-north across the northern boundary of northern domain starting at Nevada State Highway 266 (Fig. 3), where it intersects section 6. The line continues north, crosses Mount Jackson Ridge, and ends in the center of Jackson Valley, the topographic depression north of Lida Valley. A major east-west–trending gradient, increasing south to north from –9 mGal to –2 mGal (Fig. 3), is visible at the southern end of the profile. This gradient is modeled as two major south-dipping structures bounding the northern edge of the domain and reflecting vertical offsets of 160 and 130 m. Continuing north along the profile, the gravity gradient changes orientation from predominantly east-west trending to northeast trending (Fig. 3). This transition is modeled as a down-to-the-southwest extensional fault with a subsurface offset of 140 m that intersects the section obliquely and extends into Jackson Valley.

BASIN FAULT MODEL AND DISPLACEMENT BUDGET

Vertical offsets for observed and inferred faults in and around Lida Valley were combined with estimates of the regional extension direction to develop and test a 3D fault model for the basin evolution. The regional extension direction of N60°W is established for late Miocene to contemporary deformation in the region (Oldow, 1992, 2003; Oldow et al., 2008; Katopody et al., 2013) and provides a critical basis for computing a finite displacement budget for the faults bounding and dissecting the basin. Specifically, we assessed the contributions to the net slip on individual faults from the vertical components of displacement determined by offsets of topography, geologic units, and depth estimates from gravity modeling as input parameters. The vertical component of displacement was used to calculate the net slip on structures with variable but known orientation with respect to the regional extension direction. Internal consistency of our net-slip estimates was provided by the fact that several known or inferred faults are curviplanar, and as a consequence will be manifest by different vertical components of displacement for constant slip. Cumulative slip estimates were computed for all faults encountered along basin-crossing transects oriented parallel to the regional extension direction. In a system of faults where slip is conserved, similar cumulative displacement estimates along each transect are results for an indication of internal consistency within the geometric fault model for the basin.

Fault Model

The residual complete Bouguer anomaly, 3D depth inversions, and 2D forward models were combined with the regional fault map to construct the geometry of the structures underlying Lida Valley. Prominent gradients visible in the RCBA and 3D depth inversion models provide powerful controls on the orientation of major structures, which must conform to the gradient trends. The 2D forward models provide a more detailed analysis of the observed gravity and allow assessment of the number and magnitude of displacement of structures constituting the major gravity and depth gradients. We extracted the fault intersections and facing direction of structures interpreted in the 2D models as control points for the lateral continuity of structures from line to line. The projection of mapped faults and the regional gradients in depth and gravity, together with the connection of faults with similar magnitude and facing from the 2D sections, allows construction of the 3D distribution of faults within the basin.

The rectilinear geometry of Lida Valley to a first order suggests an origin as a pull-apart basin (Burchfiel and Stewart, 1966; Mann et al., 1983; Dooley and McClay, 1997) bounded by north-northeast extensional faults to the west and east and by left-oblique transcurrent faults along the north and south boundaries. Based on large-scale gradients in the RCBA and depth models, it is clear that the valley is segmented into a northern domain containing two deep subbasins that are separated from the southern domain, where the subsurface morphology is consistent with alternating horsts and grabens bounded by relatively low-magnitude structures. The intervening west-northwest boundary zone is inferred to be a distributed system of tear faults that separates the domains separating areas of localized and distributed extension. The western margin of the basin, underlying the western domain, is interpreted as a displacement relay system connecting a major down-to-the-east fault along the eastern face of Magruder Mountain and the east-dipping structures that flank exposures of Paleozoic rocks in the northwestern part of Lida Valley (Fig. 6).

The regional gradients in the RCBA and depth models clearly delineate several major structures that define the framework of the basin. The most obvious structures are the basin-bounding faults, several of which have surface expression. The inferred structures that segment the basin do not have surface expression and are solely based on our gravity models. The north and south boundaries of the basin are marked by east-west faults that are kinematically linked to north to north-northeast structures. The western third of the basin is characterized by north-northeast–trending structures and a shallow basement. This differs substantially from the eastern part of the basin, which is characterized by deep localized basins in the north and numerous shallow basins in the south separated by a tear fault with an orientation similar to the north and south structures.

The north and south boundaries of Lida Valley are defined by systems of east-west–striking faults, which, together with the regional extension direction of N60°W, predictably are dominated by left-oblique motion. The northern basin boundary is the Mount Jackson Ridge fault zone, which consists of an array of structures (Fig. 6) that trend along the southern flank of the Mount Jackson Ridge, where several fault segments are mapped. Most of the fault zone is located by an east-west–trending gradient in the RCBA and depth inversions that are slightly offset from one another at the northwestern and northeastern parts of the valley. The subsurface expression of the Mount Jackson Ridge fault zone is consistent with the projection of a mapped east-west–trending fault zone exposed in bedrock at the northwest corner of Lida Valley. The structures project from north of Lida Canyon and are tracked into the basin where they are imaged as 2 south-dipping structures that record modest vertical offsets ranging from 130 m to 160 m (Fig. 5, sections 8 and 6, west to east). These structures have a dogleg geometry and are offset along two north-northeast–striking faults mapped across Mount Jackson Ridge. East of Mount Jackson, the east-west fault zone continues along the southern flank of the ridge and projects into Stonewall Flat through a topographic depression at the northeastern margin of the valley. The southern basin boundary is simpler than the northern border and is marked by an east-west–trending gradient in the RCBA and depth inversion that represents the central strand of the Sylvania fault system as it extends across the basin from Tule Canyon (Fig. 2) east to Mount Dunfee and beyond. Only the western and eastern ends of this system are exposed and, for most of the boundary with Lida Valley, it is suballuvial and approximately located (Fig. 5, sections 1–4 and 7) with a significant down-to-the-north component of displacement ranging from 240 m to 280 m in the subsurface.

The western and eastern margins of the valley are primarily defined by the topographic expression of the bounding highlands and are marked by the East Magruder Mountain fault and by the Jackson Wash fault, respectively. Both faults trend north-northeast and dip toward the basin as they track the basin-facing edges of the bounding highlands. The sense of slip on the eastern and western basin-bounding faults is inferred to be primarily dip slip with respect to the extension direction. The western margin of the basin is dramatically controlled by the East Magruder Mountain fault, which, based on gravity modeling, is juxtaposed to a relatively shallow basin floor (Fig. 5, section 2), consistent with basalt exposures (Fig. 2). The fault flanks Magruder Mountain (Reheis and Noller, 1991) to the northwest corner of the basin, just south of Lida Valley. The eastern boundary of the basin is defined by the Jackson Wash fault, which, based on poor constrained gravity data, records a modest down-to-the-north component of ∼100 m as it flanks the eastern face of the Cottontail Ranch Hills to the northeast corner of the basin where it links to the southern strand of the Mount Jackson Ridge fault zone.

Faults in the basin are complicated and coincide with the three structural domains of the shallow western part of the basin and for the eastern part, coincide with the deep basins in the north, and low-amplitude highs and lows along the southern flank of the basin defined by the regional gradients of the RCBA and the depth inversions. Although inconvenient gaps in data coverage frustrate continuous lateral correlation of structures, the data coverage is adequate, when combined with mapped faults, to construct an internally consistent 3D fault geometry.

Two separate fault systems transfer displacement through the Lida Valley basin (Fig. 6). The eastern system extends from the eastern basin edge, underlies about two-thirds of Lida Valley, and involves the complex interaction of structures that contribute to the morphology of the northern and southern domains. The northern domain is definitively separated from the southern domain by an east-west– to west-northwest–trending gradient in the RCBA that is clearly defined both depth inversions. This gradient is modeled as a west-northwest–trending fault (Fig. 5, sections 1, 4, 5, and 7) with a modest down-to-the-north component of 100–150 m. The structural origin for this boundary is required to account for the pronounced difference is depth and character of the basement in the northern and southern domains. As such, the structural model is compatible with the gravity and associated depth models, but is not explicitly required by the potential field data.

The structural configuration of the northern domain is controlled by six north-northeast–trending faults that localize extension in two subbasins. All but the easternmost fault project south to the boundary with the southern domain but do not pass beyond the west-northwest tear fault. The easternmost fault marks the eastern flank of the basin and, although approximately located at the western margin of bedrock exposures in the Cottontail Ranch Hills, shows no evidence of being offset by the inferred tear fault. The location and orientation of the structures within the northern domain are constrained primarily by the regional gradients. The specific geometry for our model (Fig. 6) is largely controlled by the output of 2D forward models, which depict the number and magnitude of throw for specific faults. The north-northeast–trending gradient that defines the eastern boundary of the northern domain is due to two west-dipping extension faults (Fig. 5, section 4) with offsets of 200 and 250 m that gradually step displacement down into the eastern subbasin. The location of the eastern fault strand is poorly constrained by the lack of data in the eastern margin of Lida Valley and is inferred purely from the western boundary of Paleozoic units exposed at the southern basin edge, just north of Mount Dunfee. The gravity high that divides the northern domain into two subbasins is due to a major north-northeast–trending structure with a significant down-to-the-east component of ∼300 m that can be traced out of the basin (Fig. 5, section 6) and linked to an east-dipping structure mapped to the northern basin edge through Mount Jackson Ridge. This structure is in section 5 in two places as the fault parallels the section line and crosscuts the strands of the Mount Jackson Ridge fault zone at its northern end. Four north-northeast–trending faults, all with significant down-to-the-east components ranging from 250 to 570 m, are required to explain the prominent gradient that defines the western boundary of the northern domain. Three of these structures can be modeled in sections 1 and 6 (Fig. 5) with similar throws. The westernmost fault of this zone is relatively well defined and aligns with the alluvial-bedrock fault strand marking the eastern margin of Paleozoic rocks west of the northern subbasin deep. The curved nature of the easternmost fault in this north-northeast–trending zone of structures is inferred from the lateral change magnitude of vertical displacement measured on 2D profiles that cross the structure. The fault throw decreases from south to north by ∼350 m along (Fig. 5, sections 6 and 1) and requires a curvilinear geometry. In contrast, the central fault of this north-northeast–trending fault zone is best explained as projecting to the south in a curvilinear trend that parallels the orientation of a prominent regional gravity gradient (Fig. 5, sections 2 and 7). We link this structure to the central strand of the Sylvania Mountain fault system.

The structures in the southern domain of the basin are more easily distinguished and primarily constrained by the orientation of the gradients in the RCBA. At least seven north-northeast–trending, low-magnitude faults distribute displacement over a series of alternating horst and graben structures. These structures link to and do not crosscut the central strand of the Sylvania Mountain fault system forming the southern basin boundary. The eastern boundary of the southern domain is defined by the west-dipping fault that separates the basin fill sediments from a Paleozoic outcrop just north of Mount Dunfee and extends north to define the eastern boundary of the northern domain. Just northwest of Mount Dunfee, in the southern domain, an east-dipping fault is inferred based on the abrupt increase from –4 mGal to +1 mGal over a distance of ∼ 1 km in the RCBA. This position of this structure is not constrained by a 2D section line and is inferred from the juxtaposition of the RCBA transition. Three west-facing structures with vertical components ranging from 100 m to 50 m (Fig. 5, sections 1 and 2) are modeled for a modest north-northeast–trending gradient in the RCBA.

The structure of the western part of the basin is best explained as a displacement transfer system that serves as a relay (Trudgill and Cartwright, 1994) between the East Magruder Mountain fault and the structures underlying the eastern flank of Paleozoic rocks in northwest Lida Valley. The relay is composed of three faults with displacement fed from the northern strand of the Sylvania fault system. The identification of structures in the western subsurface is somewhat compromised by sparse data coverage, but due to the existence of mapped faults and the exposures of basalt in the southwestern part of the basin (Fig. 2), there is good evidence for a major fault relay system. The East Magruder Mountain fault forms the western segment of the relay and carries displacement north from the northern Sylvania Mountain fault. The displacement on the fault is estimated as the height from the top of Magruder Mountain to the basin bedrock and, based on changes in topographic expression, decreases from 2750 m to 2200 m south to north. As the elevation decreases, it is accompanied by a concomitant emergence from the basin of Paleozoic outcrops to the east. Paleozoic rocks are exposed in a triangular outcrop of bedrock south of Lida Canyon (Fig. 2) that is bounded on the west by a west-dipping fault, which continues south for ∼3.5 km based on the truncation of the western margin of an alluvial fan. The southern projection of the fault is encountered in section 3 (Fig. 5) at two locations, because the gravity line more or less parallels the structure, and is crossed in section 2 (Fig. 5). The fault is not encountered in section 7 to the south and is inferred to merge with a southern extension of one of the north-northeast–trending faults forming the western margin of the deep subbasin in the northern domain. Two small-magnitude structures are required to explain the position of basalt outcrops juxtaposed to the eastern edge of a small graben (Fig. 5, section 2) in the western domain. These faults intersect the west-dipping relay structure at their northern end and extend south, paralleling the orientation of the relay fault, to link west-northwest trend strands of the Sylvania Mountain fault system that are mapped to the basin edge.

Secondary structures were also identified at the southern end of the basin from the geophysical analysis and are consistent with projections of mapped faults in the Sylvania Mountains. Our 2D models indicate that these structures have modest displacement and do not crosscut the central strand of the Sylvania Mountain fault system. They are located to account for exposures of basalt.

Displacement Component Determination

Relative vertical offsets for Lida Valley faults with late Miocene to Pleistocene activity were evaluated to develop a displacement budget for the basin and to ensure that the proposed fault model is internally consistent. As a first step in this assessment, we established the contact between Cenozoic and pre-Cenozoic rocks as a reference datum for our reconstruction of fault displacements. The Cenozoic rocks, although displaying important variations in thickness and stratigraphy around the Lida Valley basin, for the most part are distributed as widespread units that can be traced over tens of kilometers and well beyond the bounds of our study area (Noble et al., 1964; Albers and Stewart, 1972; Weiss et al., 1993). Variations in thickness and stratigraphy typically are gradational and point to a subdued topography for the region before the onset of localized basin development on high-angle faults after ca. 8.5 Ma. We use the maximum elevation of pre-Cenozoic rocks at Magruder Mountain, and restore all structures to this reference as a means of estimating total vertical motion on faults. The vertical component of displacement for all structures was estimated from the subsurface offset of the contact between the Cenozoic and pre-Cenozoic units determined at depth from the 2D sections (Fig. 5, sections 1–8) added to the topographic difference from the basin surface to the reference datum at the approximated location of all structures. Vertical offsets of the erosional surface across faults within the highlands and bounding the basin are measured using a 10 m DEM and geospatially referenced images in Google Earth. Offsets on basin-bounding faults were restored back to the highest elevation of the Cenozoic or pre-Cenozoic reference contact in the highland adjacent to each bounding structure. Relative vertical offsets on buried interbasin structures were restored to the Cenozoic and/or pre-Cenozoic contact at depth where offset by the basin-bounding faults. Estimates of vertical displacement are considered to be a minimum due to the effects of off-fault deformation or the possibility of unrecognized fault strands.

Using the estimated vertical offset together with the strike and true dip of each structure in combination with the inferred regional extension direction (N60°W), the horizontal extension and net slip were derived for each structure. A fault dip was set at 60° for all north-northeast–trending extensional structures with strikes nearly orthogonal to the extension direction. From this condition, the magnitude of the horizontal extension was calculated for each fault using the measured vertical displacement, the acute angle between the strike and dip of the fault, and the extension direction. To calculate the horizontal extension, the vertical component of displacement must first be resolved to a horizontal plane in a direction normal to the strike of the structure. A straightforward means to compute the dip-slip component (D) of a fault as projected to a horizontal plane using the vertical component of displacement and an estimate for fault dip is: 
graphic
where HC = horizontal component, θ = dip angle of fault, and V = vertical component of displacement (throw) on the fault.
Once the vertical component is translated to a horizontal plane in a direction normal to the strike direction of the structure, horizontal extension is calculated using the value of the dip-slip component projected to the horizontal and the acute angle between the strike of the fault and the inferred extension direction. The azimuth of horizontal extension in this case is constrained to an extensional direction of N60°W (Oldow et al., 2008; Biholar, 2011; Katopody et al., 2013). The horizontal extension (HE) using the dip-slip component projected to the horizontal (calculated in Equation 1) and the acute angle between the strike of the fault and the known extension direction is: 
graphic
where DHC = dip slip projected to horizontal and Φ = angle between strike of fault and primary extension direction (ε1).
The net slip (NS) of each structure was solved for using the estimated vertical offset accompanied with the horizontal extension calculated in Equation 2, assuming that the horizontal component of slip is parallel to the extension direction. Equation 3 is used to calculated net slip using the aforementioned components. 
graphic
Net slip values calculated for structures at locations orthogonal to the extension direction of N60°W were used to determine the dip of the faults necessary to conserve slip as the structures change in orientation. Many of the faults within Lida Valley appear to be curviplanar and in some cases vary considerably along strike. Where faults change in orientation from normal to the extension direction, dips necessary to conserve slip were calculated using a derivative of the aforementioned equations to solve for θ (dip of the fault): 
graphic
where NS is held constant.

Fault Model Evaluation

The fault model developed for Lida Valley was tested for internal consistency. The cumulative horizontal displacement values for all faults along transects across the basin were calculated. The transects are aligned with the N60°W extension direction (Fig. 7). Both the net slip on faults and their horizontal displacements are computed for each fault. The aggregate totals for the horizontal components of displacement for all faults along each of the transects are compared and, for a geometrically feasible array of kinematically linked faults, should be the same. A significant discrepancy between displacement summations indicates that the geometric model is model is implausible, that the blocks accommodate significant internal deformation, or that the displacement system is not closed and has components of motion being transferred into or out of the system. The successful geometric forward models performed using the hypothesized fault geometry reduce the possibility that any displacement discrepancy is associated with block deformation.

Horizontal displacement for three transects crossing the southern, central, and northern parts of the fault system was estimated as a test of internal consistency of the model (Fig. 7). The southern transect incorporates all structures of the western relay system and several strands within the eastern system of faults and the eastern part of the Sylvania Mountain fault system, and yields a cumulative horizontal displacement of ∼1320 m. The central transect has a cumulative horizontal displacement of 1290 m and crosses all structures of the western relay system, the western segments of the faults in the northern domain, the boundary zone fault separating the northern and southern domains, most of large-magnitude faults in the southern domain, and the central strand of the Sylvania Mountain fault system. The northern transect (Fig. 7) extends from Paleozoic outcrops at the northwestern margin of the basin to the Cottontail Ranch Hills on the eastern boundary of the basin. This transect incorporates the eastern leg of the relay system at the northwest corner of the basin and all structures of the northern domain. The total horizontal extension across this transect is estimated as 1400 m.

The southern and central transects yield remarkable internal consistency in horizontal displacement but are ∼100 m less than the cumulative displacement along the northern transect. A maximum discrepancy of only 100 m between transect displacements is, in light of the uncertainties associated with depth models from gravity and the restoration datum used, virtually indistinguishable and supports the internal consistency of the fault model. The northern transect incorporates two structures that crosscut the Mount Jackson Ridge fault system along the northern flank of the basin (Fig. 7). The magnitudes of net slip and horizontal components of displacement for each of these structures is greater south of the Mount Jackson Ridge fault zone than for the fault segments farther north. The implication is that truncation of the Mount Jackson Ridge fault system occurred late in the displacement history, with strands of basin-bounding faults beneath Lida Valley transferring displacement out of the fault network. The contribution of horizontal displacement for the two faults north of the basin is ∼110 m, which must be removed from the transect total. This yields net horizontal displacements that are comparable within ∼10% across the entire basin fault system.

We calculated displacement budgets on individual fault systems within the basin fault array as a means of assessing the source of model discrepancies. The displacement on segments of the western relay were assessed, as were the cumulative displacements for the northern and southern domain structures. The model for the western relay system yields horizontal components of displacement of 765 m in the south and 670 m in the north (Fig. 7) and suggests that the displacement budget at the northern end of the system, where data coverage is poor, may be underestimated. The displacement budgets for the structures in northern and southern domains of the basin, measured along three transects, are 905, 980, and 1100 from south to north (Fig. 7) and represent a maximum discrepancy of 18%. Although discrepancies are more pronounced in individual systems of faults, the overall result serves as additional validation for the fault model.

DISCUSSION AND CONCLUSIONS

The Lida Valley basin is bound and crosscut by a rectilinear array of faults that emanate from the Sylvania Mountain fault system and transfer displacement to the Mount Jackson Ridge fault zone. The Lida Valley is a pull-apart structural system with extension localized on north-northeast faults and transcurrent displacement on east-west faults. The combination of mapped faults and subsurface faults associated by gravity data, however, indicate an internal geometry that is substantially more complex that most pull-apart basin systems. The system of faults within the basin is composed of fault relays and transfer fault systems that segment the basin into three structural domains (Fig. 4C). Faults within each structural domain exhibit a tendency to curve along strike and to merge with other faults, resulting in an intricate array of kinematically linked structures that transfer displacement through the basin.

In the western domain, two kinematically linked fault systems transferred motion between the south and north bounding faults (Fig. 6). The westernmost fault system is a relay (Trudgill and Cartwright, 1994; Walsh et al., 1999) consisting of three fault strands that link the east-west–trending Sylvania Mountain and Mount Jackson Ridge faults. The relay appears as a Z in map view. The western leg of the relay transfers displacement from the Sylvania Mountain fault at the southwest corner of the basin to the north along the east-dipping East Magruder Mountain fault. Displacement gradually decreases along the East Magruder Mountain fault to the north and is transferred to a southwest-dipping fault bounding Paleozoic basement exposed to the east (Fig. 2). Displacement on the west-dipping fault decreases to the south and in turn is transferred to a north-northeast–striking, east-dipping fault that forms the western margin of the deepest subbasin in the basin. Displacement on the eastern leg of the relay fault is passed to the east-west–striking Mount Jackson Ridge fault zone along the northern flank of the basin. The eastern fault system in this area is a curved transfer fault that carries displacement from the central Sylvania fault system north to the Mount Jackson Ridge fault zone.

The geometry of the eastern part of the basin differs substantially from that of the west. The eastern part of the basin is segmented into structural domains characterized by two major subbasins to the north and a series of small basins to the south separated by a west-northwest–trending transfer fault (Fig. 6). The magnitude of extension within the structural domains is the same but is localized in two subbasins in the north and distributed across a series of horst and graben structures in the south. The west-northwest fault zone separating the domains serves as a tear fault.

Cumulative displacement on the array of faults of the Lida Valley basin system is modest and provides an estimate of magnitude of displacement transferred from the FC-FLV fault to the central Walker Lane on the Sylvania fault since the late Miocene. The net vertical displacement accumulated on Lida Valley high-angle faults is ∼2.3–2.5 km; this yields a net slip of 2.6–2.9 km for the system. The horizontal component of displacement on the fault array ranges between 1.3 and 1.4 km, and represents only ∼10% of the left-lateral displacement recorded for the Sylvania Mountain fault system. The implication is that most of lateral displacement on the Sylvania Mountain fault system either predates formation of the Lida Valley basin or is accommodated by a different mechanism, such as vertical axis rotation of rigid blocks. Our analysis provides an important constraint on the displacement budget for part of the network of faults underlying the southern Walker Lane.

This research was supported by National Science Foundation grant EAR-0948552 and funds provided by Pioneer Natural Resources, Inc.