Geologic mapping, 40Ar/39Ar dating, and whole-rock geochemical analyses were employed to establish the Cenozoic stratigraphy, geometry and timing of normal faults, and the magnitude of extension in the vicinity of Golconda Canyon in the southern Tobin Range. The Golconda Canyon area is near the westernmost extent of a major east-west–trending paleovalley that likely predates Basin and Range extension in that region. Latest Eocene–Miocene volcanic rocks infill and overtop the paleovalley.

Four phases of extensional normal faulting have been identified in the Golconda Canyon area. The earliest phase of normal faulting was minor and took place during the early Oligocene in the form of northwest– and northeast–striking, west-dipping faults. This faulting resulted in as much as ~5°–10° of tilting in the western part of the Tobin Range, and occurred coeval with andesitic volcanism. A second phase of faulting consisted of a major northwest-striking fault and associated northeast-striking faults, both of which dip west. These faults may have begun movement as early as the late Oligocene(?), with much displacement in the middle Miocene (pre–14.1 Ma). The faults produced syntectonic basins filled with landslide breccia and megabreccia as well as coarse clastic sedimentary rocks. The next phase may represent an eastward progressive younging of the earlier phase, because it is documented largely in the middle and late Miocene record of the eastern part of the range. These north-striking faults are mostly west dipping, but some are east dipping, and they accommodated ~25°, and locally 30°, of eastward tilting. Younger moderately to steeply dipping active faults bound the eastern and western margins of the Tobin Range, and include the western fault that ruptured during the 1915 Pleasant Valley earthquake.

Cumulative normal faulting has resulted in an overall ~25°–30° eastward tilt of the Tertiary stratified rocks in the Golconda Canyon area. The timing and magnitude of this faulting and tilting vary, with all areas showing the greatest magnitude of faulting and higher extension rates associated with the two sets of Miocene faults. The western part of the range displays evidence of early Oligocene faulting associated with andesitic magmatism, whereas these faults are not exposed in the eastern part of the range, where middle Miocene and younger tectonism produced slightly greater total stratal tilts.

Palinspastic restorations of cross sections and calculations based on fault and rock unit dips suggest that the southern Tobin Range in central Nevada has undergone a minimum of ~50% east-west crustal extension since 34 Ma. This amount of extension and the initiation of normal faulting ca. 33 Ma in the Tobin Range are consistent with the westward decrease in the age and magnitude of extension at this latitude in the Basin and Range Province.


In the northern Basin and Range Province, extension and magmatism began in the late Eocene and have generally migrated southward and westward (Seedorff, 1991). Numerous studies have been conducted to estimate the amount and timing of extension via Cenozoic normal faulting at various locations in the Nevada part of the Basin and Range Province (i.e., Proffett, 1977; Wernicke et al., 1988; Smith, 1992; Hudson et al., 2000). The southern Tobin Range in Pershing County, Nevada (Fig. 1), is of considerable interest for understanding extensional structural geology because it is near the apparent transition between more extended terrane (~100%) to the south and east (Proffett, 1977; Smith et al., 1991; Hudson et al., 2000), and generally less extended terrane (~15%) to the north and west (Lerch et al., 2004; Colgan et al., 2006a, 2006b). Despite these different estimates of extension across this apparent boundary, geophysical data suggest that the crust is relatively thin on both sides of the transition (Klemperer et al., 1986).

We have employed geologic mapping, descriptions of volcanic and syntectonic sedimentary deposits, and radiometric dating of volcanic units to better understand the nature of late Cenozoic normal faulting in the Tobin Range. With these data we endeavor to (1) describe the geometry of normal faults, (2) establish the relative and absolute ages of faulting, (3) estimate the total amount of extension, and (4) place the Tobin Range in the context of the aforementioned extensional boundary.

Regional Geology

Pre-Cenozoic Geology

In northern Nevada, Proterozoic continental rifting produced a west-facing passive continental margin, and subsequently in the Paleozoic at the longitude of the Tobin Range, deep-marine pillow basalts and chert-argillite sequences were deposited (Jones and Jones, 1991). In the late Paleozoic, the Roberts Mountain and younger Golconda faults thrust the marine sequences eastward onto the continental margin (cf. Silberling and Roberts, 1962; Snyder and Brueckner, 1983). Thick sequences of marine to subareal sedimentary and volcanic rocks were deposited in the Triassic and Early Jurassic. A belt of Mesozoic thrust faults is west of the Tobin Range, as is the zone of Middle Jurassic and Cretaceous Cordilleran batholiths.

Cenozoic Geology

In the north-central Nevada part of the Basin and Range Province, approximately east-west–directed extension accommodated by north-striking normal faults has resulted in the development of tilted range-forming fault blocks separated by alluvial basins (Figs. 1 and 2). Normal faulting and extension in Nevada have migrated with time, such that the older faulting episodes are found in northeastern Nevada, and with time the zone of maximum faulting and extension rates has migrated south, west, and east (e.g., Christiansen and McKee, 1978; Gans et al., 1989; Seedorff, 1991; Wernicke, 1992; Fig. 2). Extension has been broadly associated with magmatism; late Eocene (35–40 Ma) extension in northeastern Nevada is closely associated with the initiation of dacitic magmatism (cf. Gans et al., 1989; Feeley and Grunder, 1991), Carlin-type and epithermal gold deposits, and porphyry copper-gold mineralization (Ressel and Henry, 2006; Castor et al., 2003). Oligocene extensional faulting documented for the Toiyabe Range, Stillwater Range, and Atlanta district (Fig. 2) are associated in time and position with caldera complexes that are the source of large silicic ignimbrite sheets extending in a belt across central Nevada.

In the Walker Lane belt between Las Vegas and Reno, extension, normal faulting, and tilting began locally in the late Oligocene, whereas widespread patchy zones of extension associated with andesitic magmatism began ca. 15 Ma (cf. Proffett, 1977; Dilles and Gans, 1995; Surpless et al., 2002). In the Walker Lane belt, northwest-striking dextral strike-slip faults have been active over the same period as normal faulting (cf. Dilles and Gans, 1995). Similarly, Colgan et al. (2008) documented extension in north-central Nevada.

In the past 15 m.y., the locus of faulting and extension has moved westward to the edge of the modern Sierra Nevada block and propagated northwestward (cf. Henry et al., 2004, 2007; Carmichael et al., 2006; Colgan et al., 2006b). At present the most active faulting in Nevada is concentrated along the Sierra Nevada front, in the Walker Lane belt and in the Central Nevada seismic zone, which extends northward from the Walker Lane through the western margin of the Tobin Range. The seismically active Central Nevada seismic zone has undergone a number of historical earthquakes (Fig. 1A), including four large surface-rupturing earthquakes of magnitude > 6.5. From north to south, they are the 1915 Pleasant Valley earthquake (Wallace, 1984a), which ruptured along the western margin of the Tobin Range, the 1954 Dixie Valley and Fairview Peak earthquakes (Slemmons, 1957), and the 1932 Cedar Mountain earthquake (Gianella and Callagan, 1933, 1936).

The elevated heat flow associated with volcanism may have thermally weakened the crust and aided extension in the Basin and Range Province (e.g., Gans et al., 1989). Alternatively, Best and Christiansen (1991) suggested that periods of increased volcanism do not coincide with times of extension. Despite continued faulting and nonhomogeneous extension of the upper crust, each part of the Basin and Range Province remains in isostatic equilibrium through ductile flow of the lower crust and upper mantle (e.g., Gans, 1987; Block and Royden, 1990; Wernicke, 1990; McKenzie et al., 2000).

Today, the greatest strain rates in the Basin and Range Province are generally along the eastern and western margins of the province (DePolo et al., 1991). Recurrence intervals for faults in the province are estimated to be several thousands of years to more than 100 k.y., and estimated average slip rates for large, range-bounding faults are between 0.01 and 1.0 mm/yr (Wallace, 1984b, 1987; Wesnousky et al., 2005).

Major characteristics of the Basin and Range Province include high average elevation, high heat flow, and thin crust (Stewart, 1978). Heat flow in the Province ranges from 60 mW m−2 in the Eureka low, in central Nevada, to 110 mW m−2 in the Battle Mountain high, in which the Tobin Range is situated (Blackwell et al., 1991). Catchings and Mooney (1991) estimated crustal thicknesses of 32–36 km with the use of seismic reflection profiles at lat 40° N.


Geologic Mapping

We conducted 10 weeks of geologic mapping at a scale of 1:12,000 during the summers of 2004 and 2005 on parts of the Needle Peak and Jersey Summit 7.5′ quadrangles and used U.S. Geological Survey topographic maps and digital orthophotographs as base maps. These geologic data were compiled together with previous data to produce a geologic map of the southern Tobin Range (scale 1:100,000; Fig. 3). The detailed geologic map and supplemental cross sections are presented in Plate 1 (scale 1:24,000). Cross-sectional drawings, traverses up well-exposed rock exposures, thin sections, major and trace element analyses, and age determinations serve as the basis for construction of stratigraphic relationships.

40Ar/39Ar Ages

Five samples of volcanic rocks that display little alteration were chosen from key positions in the Cenozoic stratigraphic section. Pure mineral separates of plagioclase (n = 2) and sanidine (n = 2) together with one cleaned whole-rock sample were analyzed via the 40Ar/39Ar step-heating method in the Oregon State University laboratories.

Details of 40Ar/39Ar Age Methodology

All samples were cleaned of weathered surfaces, crushed, sieved, and cleaned in deionized water and dilute HCl in an ultrasonic bath. Use of a Frantz isodynamic separator and hand-picking under a binocular microscope assured 100% purity of mineral separates. Samples TR-9 and TR-81 were plagioclase separates, samples TR05-21 and TR05-26 were sanidine separates, and TR-77 was a crushed whole-rock sample. For full procedural details, see Gonsior (2006).

Analytical procedures followed methods described in Duncan and Keller (2004). Samples were wrapped in Cu foil, packed with Fish Canyon Tuff biotite flux monitor standard (28.03 ± 0.18 Ma; Renne et al., 1994), placed into a sealed silica tube, and irradiated for 6 h at Oregon State University's TRIGA experimental reactor. Ages were determined at the Oregon State 40Ar/39Ar Geochronology Lab using two gas extraction methods. Four samples were incrementally step-heated with a double vacuum Heine low-blank resistance furnace with a Ta-Nb crucible. Gas was incrementally heated at 10–12 steps between 600 and 1400 °C, cleaned with Zr-Al and Zr-V-Fe getters, and then masses 36, 37, 38, 39, and 40 were measured sequentially on a Mass Analyser Products 215-50 gas mass spectrometer.

The fifth sample (TR05-21, sanidine) was step-heated by laser to fusion due to its low mass (49 mg) using a Merchantek 30 watt CO2 continuous-fire laser. The 50 μm laser spot step-heated the sample from 400 to 1500 °C as temperature was monitored with API Micro Probe model MSP-200G. Software from Koppers (2002) was used to reduce data, determine isotopic ratios, and calculate ages using the known age of the flux monitor.

The following criteria, developed by Dalrymple and Lanphere (1974), Duncan et al. (1997), Tegnar and Duncan (1999), and Frey et al. (2004), were used to determine accepted closure ages: (1) a plateau is established with a minimum of three contiguous heating steps within the limits of a 2σ error, and represents ≥50% of total 39Ar released (the steps are used to calculate the weighted mean plateau age); (2) a mean square of weighted deviates (MSWD) analysis of the weighted mean plateau age is >1 but <2.5; (3) the plateau age is concordant with the isochron age and; (4) the initial 40Ar/36Ar ratio calculated from the isochron is within a 2σ error limit of the 40Ar/36Ar ratio for atmospheric argon.

Of the five samples analyzed, four met the aforementioned criteria and yielded the robust plateau ages used herein (Table 1; Fig. 4). Sample TR-81 did not develop an age plateau, but rather a U-shaped pattern of argon release, indicating excess argon. The isochron yields an unacceptable MSWD, <1. Five steps with 72% of the 39Ar yield an age of 40.38 ± 5.7 Ma, but we here select the youngest age step of 36.58 ± 0.29 Ma as the maximum age estimate and the inverse isochron age of 24.63 ± 3.61 Ma as the minimum possible age. The full data are available in Supplemental Table 11, and sample locations are show in Plate 1 and Supplemental Table 2 (see footnote 1).

Geochemistry of Volcanic Rocks

Whole-rock geochemical analyses were obtained on 12 samples of rhyolitic ignimbrites and 2 samples of andesitic rocks to characterize volcanic units and aid in correlation. Samples were taken of fresh rocks, except for rhyolite ignimbrite clasts in the landslide breccia unit, that were all affected by weak hydrothermal alteration. Major elements were analyzed by X-ray fluorescence following lithium borate fusion, and trace elements by combined inductively coupled plasma–mass spectroscopy (ICP-MS) and ICP–atomic-emission acid spectroscopy following HF-HNO3-HClO4 digestion and HCl leaching by Chemex Laboratories. In the latter method, refractory phases such as zircon and apatite may sometimes be incompletely dissolved. The data are presented in Table 2 (normalized to 100 wt% anhydrous oxides) and summarized in Figure 5, and detection limits for respective methods are presented in Supplemental Tables 3, 4, 5, 6, and 7 (see footnote 1).


Pre-Cenozoic Rocks

The oldest rocks in the study area belong to the Pennsylvanian–Permian Havallah Formation, a thick sequence of quartzite, chert, argillite, limestone, and rare greenstone (Muller et al., 1951; Ferguson et al., 1952). Also in the vicinity, but not in the map area, is the Pennsylvanian–Permian Pumpernickel Formation, a thick sequence of greenstone, chert, and argillite that underlies the Havallah Formation (Muller et al., 1951). The Havallah Formation crops out along the northern edge of the field area, in the vicinity of Needle Peak, and underlies much of the central Tobin Range to the north. Regionally, these rock units represent the deep-marine sequences emplaced along the latest Paleozoic Golconda thrust (Silberling and Roberts, 1962).

The Lower Triassic Koipato Group overlies the Havallah Sequence and consists of four units: the Limerick Greenstone (oldest), the Rochester Rhyolite, the Weaver Rhyolite, and the China Mountain Formation (youngest). In the Golconda Canyon area, lithologies include tuffaceous sandstone, rhyolite- and chert-pebble conglomerate, and lesser amounts of rhyolite, and are assigned to the China Mountain Formation. The age of the Koipato Group is 225 ± 30 Ma on the basis of a fission-track age on zircons from Rochester Rhyolite samples collected in Pleasant Valley (McKee and Burke, 1972).

Burke (1973) provided detailed descriptions of the pre-Tertiary stratigraphy in the southern Tobin Range near and south of Golconda Canyon. The Natchez Pass Formation of the Middle and Upper Triassic Star Peak Group and the Osobb Formation of the Upper Triassic Auld Lang Syne Group overlie the Koipato Group. The Natchez Pass Formation is composed of dark gray to brown massive limestone and dolomite, whereas the Osobb Formation is predominantly sandstone.

A granodiorite of unknown age intrudes the Havallah Formation and the Koipato Group. On the Pershing County geologic map, Tatlock et al. (1977) assigned this small intrusion a Jurassic age. Similar granodiorite intrusions in central Nevada have yielded Jurassic K-Ar ages (McKee et al., 1971), but Permian, Cretaceous, and Tertiary granodiorites have also been mapped in Pershing County (Tatlock et al., 1977).

Cenozoic Rocks

Cenozoic Paleovalley

An early Tertiary east-west–trending paleovalley formed near Golconda Canyon and was infilled, and overtopped by Tertiary volcanic and sedimentary units (Plate 1 and Fig. 6). A second, less deep paleovalley with Caetano Tuff as the basal unit is exposed 10 km to the south in the Tobin Range (Burke, 1973; Fig. 3). Masursky (1960) and Burke and McKee (1979) originally described the Golconda Canyon paleovalley as a volcanotectonic trough that developed during the late Eocene and extended until the Miocene. Near the mouth of Golconda Canyon, the local presence of gravels and conglomerates in the canyon at the base of the Tertiary section and the occurrence of vitrophyres where Tertiary ignimbrites banked into the pre-Tertiary units suggest that the genesis of the valley feature was via erosion and not directly by faulting and magmatism. The northern margin of the Golconda Canyon paleovalley is bounded along much of its length by a series of west- and northwest-striking normal faults that are Miocene age based on offset of the Fish Creek Mountains Tuff, the pumiceous rhyolite, and syntectonic sedimentary rocks and breccia deposits. The younger faults obscure much of the original geometry of the late Eocene paleovalley on its north side. Nonetheless, the earliest Oligocene Caetano Tuff, the overlying andesitic eruptive sequence, the 25 Ma pumiceous rhyolite ignimbrite, and the base of the Miocene syntectonic sedimentary rocks and breccia deposits are all deposited on top of the Triassic basement rocks on the north side of the paleovalley near the mouth of Bushee Canyon in the northwest part of the map area (Plate 1 and Fig. 6). In this area, the rapid pinch-out of these units to the north within 2–3 km of the paleovalley center indicates that this margin had high relief. The southern margin of the paleovalley had relatively low relief, as indicated by the gradual pinch-out of the Caetano Tuff southward (Fig. 3). The two margins originally trended northwest to west-northwest, on the basis of the lineation formed by the rock units pinching out against the valley walls. At least 1000 m of late Eocene–early Oligocene volcanic rocks filled the paleovalley, the width of which was ~6–7 km, based on the distribution of the valley-confined Caetano Tuff (Fig. 3). Normal faulting was ongoing when early Oligocene volcanic rocks filled the paleovalley (see following) but was small in magnitude, and the documented faults strike northeast at a large angle to the paleovalley trend. Therefore, the early Oligocene paleovalley does not appear to be principally controlled by faulting.

During the early and middle Miocene, the northern margin of the paleovalley was largely bounded by the west- and northwest-striking normal faults that have syntectonic sedimentary and breccia deposits in their hanging walls (south) that thin to the south. These relationships support the hypothesis that in the Miocene, normal faulting produced a half-graben that partly controlled the paleovalley geometry.

Older Basalts (Tob)

A thick sequence of ignimbrites, lava flows, and sedimentary rocks unconformably overlie the pre-Cenozoic rocks. The lowest units, the Eocene basalts and Caetano Tuff, are confined to the Eocene paleovalley described above. Figure 7 illustrates the stratigraphic section and relationships of the Cenozoic rock units in schematic form. The oldest unit filling the paleovalley is a series of as many as 25 basalt lava flows composing a section as thick as 200 m. The basaltic lavas have only been observed at the bottom of the paleovalley near the mouths of Golconda and Bushee Canyons, where they locally overlie a few meters of basal conglomerate and sandstone.

Caetano Tuff (Tct)

The Caetano Tuff overlies the basalt and consists of moderately crystal rich, poorly to moderately welded biotite-bearing rhyolite ignimbrite. The Caetano is as thick as 250 m and is informally divided into a poorly welded lower unit and a moderately welded upper unit (Fig. 7). The Caetano Tuff is late Eocene–early Oligocene, based on numerous age determinations ranging between 32 and 34.5 Ma (Naeser and McKee, 1970; John et al., 2003). John and Henry (2005) and John et al. (2008) proposed that there are two distinct ages and therefore two eruptive units previously considered to be the Caetano Tuff. The younger unit is designated the Caetano Tuff and has yielded sanidine 40Ar/39Ar ages of 33.84–33.77 Ma for exposures of intracaldera ignimbrite in the Shoshone Range, ~50 km east of the Tobin Range. The older unit is now designated the Tuff of Cove Mine and is defined by ages between 34.5 and 34.2 Ma for samples north of the caldera. John et al. (2008) obtained an 40Ar/39Ar age of 33.75 ± 0.06 Ma for sanidine from the Caetano Tuff collected by us from Golconda Canyon. Therefore, the Tobin Range and Shoshone Range caldera samples of the Caetano Tuff have the same age, and are considered latest Eocene based on the 33.7 Ma age for the Eocene-Oligocene boundary (Palmer and Geissman, 1999), but could also be considered early Oligocene based on the proposed 37.9 Ma boundary (Gradstein et al., 2004). The geochemical composition of four samples from Golconda Canyon is very similar to the composition of intracaldera samples in the Shoshone Range (John and Henry, 2005; John et al., 2008). The Tobin Range samples have 72–76 wt% silica and a distinctive trace element geochemical composition including high Rb (200 ppm) and Ba (200–1500 ppm) and relatively low Nb (15 ppm), Y (20 ppm), and Sr (150–350 ppm; Fig. 5). On the basis of age and chemical composition, the Caetano Tuff outflow sheet of the Tobin Range is correlated with the intracaldera and extracaldera exposures of the Shoshone Range (John et al., 2008). The Caetano Tuff in Golconda Canyon represents the westernmost and farthest traveled outflow sheets from the Caetano caldera.

Biotite Rhyolite Ignimbrite (Trib)

A moderately welded, crystal-rich, rhyolite ignimbrite with abundant biotite (5 vol%) and moderately abundant pumice is as thick as 150 m and occurs as rare, isolated exposures in the field area where its stratigraphic position is uncertain. In one exposure the biotite tuff is overlain by andesite eruptive rocks, but in two other exposures it either overlies the base of the andesite sequence or is faulted against these andesites. An 40Ar/39Ar age on sanidine of 33.28 ± 0.22 Ma (Table 1) is distinctly younger than the age of the Caetano Tuff from Golconda Canyon at the 95% confidence level, but is analytically equivalent within error to the 33.03 ± 0.25 Ma age of an andesite ignimbrite near the base of the andesite sequence. Therefore, the stratigraphic position of the biotite tuff remains uncertain, but is here placed above the Caetano Tuff and below the andesite sequence.

The geochemical composition of two samples of the biotite rhyolite ignimbrite (73–77 wt% silica but low alkalies) is within the range of the Caetano Tuff from the Tobin Range. The close match of the trace element suites for these two sets of rock samples (Plate 1A–1E, 1H) supports our hypothesis that the biotite rhyolite may represent an ignimbrite eruption from the Caetano caldera magmatic system; however, no similar unit has been noted within the Caetano caldera (John et al., 2008). Therefore, the correlation of the biotite rhyolite and its source remain uncertain, but based on its age, stratigraphic position, petrology, and geochemistry, this tuff could either be a third and youngest, and otherwise unrecognized, extracaldera ignimbrite from the Caetano caldera or derived from a source not currently recognized.

Andesite Eruptive Sequence (Tai, Tal, Tas, Tau)

A thick sequence of as much as 700 m of andesitic rocks overlies the Caetano Tuff in most exposures. The general stratigraphy near Golconda Canyon consists of a basal series of biotitepyroxene phyric ignimbrites (Tai); a thick series of pyroxenephyric andesitic lava flows and flow-breccias with intercalated lahars (Tal); and locally at the top of the section, a series of volcaniclastic sandstones and conglomerates (Tas). The ignimbrites form a series of thin (<20 m), poorly to moderately welded units intercalated with sedimentary rocks and lavas (Fig. 8A). In the lava sequence, blocky lava flows 5–20 m thick have distinctive 1-cm-long augite phenocrysts. At least three platy-jointed lava flows at the top of the sequence are prominent ridge formers and are characteristically distinct due to large (1 cm) hornblende phenocrysts or abundant vesicles. Minor amounts of landslide breccia containing rhyolite ignimbrite fragments are intercalated in both the andesite ignimbrites and in the lower part of the pyroxenephyric lava flows.

Sample TR-9 yielded an 40Ar/39Ar plateau age of 33.03 ± 0.25 Ma on plagioclase from a vitro-phyric ignimbrite near the base of the sequence (Table 1). A plagioclase from a sample (TR-81) of hornblende phyric andesite lava near the top of the section did not yield a plateau age, and we interpret a maximum age of ca. 36.6 Ma and a minimum age of 24.6 ± 3.6 Ma.

A geochemical analysis of an andesite ignimbrite at the base of the section contains 61 wt% silica, whereas a hornblende-phyric lava near the top of the section contains 64 wt% silica, and supports increasing silica content going upsection. These andesitic rocks represent a high-K (1–3 wt% K2O) suite with high Ba (~2000 ppm), Y (25 ppm), Sr (1000–1500 ppm), and La (41–54 ppm) similar to other large ion lithophile element– and light rare earth element–enriched andesites of the eastern Great Basin.

Andesitic and Dacitic Dikes (Taid, Tdi)

Numerous dacitic and andesitic dikes are associated with the andesite eruptive sequence based on composition and relative age, and many are apparently feeders to lavas; however, some dikes clearly cut and postdate the andesite lavas. The andesite dikes mostly crop out in the southern Tobin Range south of Golconda Canyon, where they intrude the earliest generation of normal faults (Fig. 3; Burke, 1973). A Miocene northwest-striking dacitic dike in the southern part of the study area is emplaced along a fault that juxtaposes Oligocene andesite lavas against the base of the overlying Miocene sedimentary rocks.

Pumiceous Rhyolite Ignimbrite (Trip)

Overlying the andesite sequence is a pumice-rich rhyolite ignimbrite. This gray to brown unwelded rhyolite is crystal poor (10 vol%), moderately lithic rich, and contains anorthoclase, sparse plagioclase, sparse sanidine(?), and as much as 30% pumice. An 40Ar/39Ar age determination on anorthoclase from a sample near the top of the unit yielded a plateau age of 24.95 ± 0.17 Ma (Table 1). This age constrains the upper age limit of the andesite extrusive rocks.

This pumice-rich rhyolite unit may be correlative to the fourth cooling unit of the Bates Mountain Tuff of central and eastern Nevada, which is likewise correlative to the Nine Hill Tuff in western Nevada and eastern California, but the correlation remains uncertain (see following review). Sargent and McKee (1969) described the Bates Mountain Tuff as being a crystal-poor ash-flow tuff with phenocrysts of sanidine, quartz, and lesser amounts of plagioclase, with varying degrees of welding. The 40Ar/39Ar age determinations by Deino (1989) on samples from both the Nine Hill Tuff and the Bates Mountain Tuff have produced a mean age of 25.11 ± 0.017 Ma that is within 2σ error of the age determined here. The high Ca/K ratio of 6 determined by our 40Ar/39Ar analysis is distinctive of the Nine Hill Tuff and peralkaline tuffs of the McDermitt caldera (C.D. Henry, 2007, personal commun.). Whole-rock trace element geochemical analyses compiled by Deino (1985) are generally similar to the compositions of the pumice-rich rhyolite in the Tobin Range, but differ in a few key characteristics (Tables 2 and 3; Fig. 5). The Nine Hill Tuff is characterized by a distinctive trace element suite that includes elevated Nb (30 ppm) and Zr (400 ppm). The Tobin Range samples (71–74 wt% silica) are similar but contain slightly lower high field strength elemental (HFSE) concentrations, with 26 ppm Nb and 190 ppm Zr. The lower Zr and other HFSE contents of the Tobin Range samples may be due to incomplete digestion of zircon and other refractory minerals in the four-acid digestion method employed prior to ICP-MS analysis. The pumiceous rhyolite of the Tobin Range, however, contains much higher Sr (80–100 ppm), Rb (200 ppm), and Ba (800 ppm) compared to the Nine Hill Tuff (20 ppm Sr, 40 ppm Rb, and 200 ppm Ba; Fig. 5F). These significantly different trace element contents suggest that the pumiceous rhyolite tuff, while being similar in age and mineralogy to the Nine Hill Tuff, may represent a different eruption. Based on the geochemical mismatch and slightly different 40Ar/39Ar ages, the pumice-rich rhyolite ignimbrite is not equivalent to the Nine Hill and Bates Mountain Tuffs and may represent a slightly younger alkali rhyolite tuff.

Fish Creek Mountains Tuff (Tfc)

Elsewhere in the southern Tobin Range, the Fish Creek Mountains Tuff, a quartz- and sanidine-rich, lithic-poor, moderately to strongly welded rhyolite ignimbrite, overlies the andesite sequence (Fig. 3). In the exposures in the detailed map area east of Golconda Canyon, the ignimbrite is as much as ~15 m thick, but landslide blocks in the sedimentary breccia unit (Tsx) are as thick as 30 m (Fig. 8B). A few ridge-forming exposures of Fish Creek Mountains Tuff can be traced a few hundred meters with consistent attitudes of compaction foliation, whereas many other exposures are brecciated, crop out as slope formers, and contain diverse attitudes of compaction foliation. Whether the latter are in place of original deposition is debatable, and is further discussed in the sedimentary rock and rhyolite landslide breccia unit description. The type locality and source area of this ash-flow tuff is a caldera complex in the Fish Creek Mountains (Burke and McKee, 1979), 10 km east of the southern Tobin Range. The Fish Creek Mountains Tuff in the type location has an age of 24.6 ± 1.3 Ma determined from fission-track dating on zircon (McKee et al., 1971). An 40Ar/39Ar age of 24.72 ± 0.05 Ma has been obtained for sanidine from the Fish Creek Mountains Tuff (John et al., 2008).

In the type locality of the Fish Creek Mountains Range (Fig. 1) the Fish Creek Mountains Tuff has two cooling units. Based on lithology and geochemistry, exposures in the southern Tobin Range correlate with the upper cooling unit of the type locality. The lower unit is poorly to moderately welded, poorly to moderately crystal rich, lithic rich, and grades upward into the upper cooling unit, which is strongly welded, crystal rich, and lithic poor.

Three samples from the rhyolite landslide breccia, here interpreted to be derived from the Fish Creek Mountains Tuff, have 73–75 wt% silica and contain low Sr (45–95 ppm) and elevated Zr (100–200 ppm); Zr content is positively correlated with Hf, elevated Nb (25–56 ppm), Y (40–65 ppm), Ta, and Th (Fig. 5, unit Trx). These trace element contents are distinctive, and differ from the Caetano and pumiceous tuffs of the Tobin Range. No published geochemical data for immobile elements are available for the Fish Creek Mountains Tuff, although McKee (1970) inferred whole-rock values for Sr and Rb from known values of Sr, Rb, and K in sanidine phenocrysts. McKee's (1970) values for Sr (65 ppm) and Rb (265 ppm) are similar to samples from the rhyolite landslide breccias analyzed here (Sr = 47–93 ppm, Rb = 199–440 ppm; Fig. 5H).

Syntectonic Sedimentary Rocks and Breccia (Tsc, Trx, Tlx)

A map unit of sedimentary rock (Tsc) and rhyolite-rich landslide breccia (Trx) overlies the Fish Creek Mountains Tuff and the pumiceous rhyolite tuff (Nine Hill Tuff?). Two distinct rhyolitic lithologies are evident and dominate the clasts present in the landslide breccia in the southern Tobin Range. One lithology is a red, crystal-rich, strongly welded, lithic- and pumice-poor rhyolite ignimbrite, similar to in situ exposures of Fish Creek Mountains Tuff in the Tobin Range. The other rock type is a hydrothermally altered, porous, whitish, poorly welded lithic- and pumice-rich, crystal-poor, rhyolite ignimbrite that may be correlative to the lithic- and pumice-rich lower cooling unit of the Fish Creek Mountains Tuff. No exposures of this rhyolitic ignimbrite have been found in situ in the Tobin Range. Also present in the landslide breccia are local blocks of limestone derived from the Natchez Pass Formation (Tlx).

The sandstones and conglomerates associated with the landslide breccia are usually poorly exposed in areas of subdued topography with abundant andesite and rhyolite cobbles and boulders within clay-rich soil (Fig. 9). The sandstones are typically poorly sorted, whereas the conglomerates are dominated with clasts of rhyolite and andesite, with lesser amounts of material derived from neighboring pre-Tertiary rocks. These sedimentary rocks are generally localized in the hanging wall of a major northwest-trending normal fault and appear considerably thicker in the northwest part of the study area (Fig. 9).

This rock unit has been previously mapped on both the geologic maps of Pershing County (Tatlock et al., 1977) and the southern Tobin Range (Burke, 1973) as in situ exposures of Fish Creek Mountains Tuff and as sandstone and conglomerate associated with the andesite sequence. Field exposures examined in this study best support an interpretation that most Fish Creek Mountains Tuff exposures in the Tobin Range are landslide breccia deposits and have been displaced from their original depositional position. The exposures of the deposits consist primarily of unconsolidated clay-rich volcaniclastic materials containing angular cobbles, boulders, or large blocks to 7 m in diameter (Figs. 8B–8D). From one block to the next, orientations of fiammé vary in a random pattern and indicate rotation of blocks relative to original depositional position. Locally, 0.5 km west of the Tobin Mine, semi-intact blocks to 200 m long of Fish Creek Mountains Tuff with relatively intact internal stratigraphy (Fig. 8B), but significant block rotation as shown by compaction foliation dips varying by 90°, are interpreted as mega-breccia landslide blocks (Fig. 8C). These large blocks in the landslide deposits are in the hanging wall (south side) of a northwest- striking normal fault that extends across the range or near the base of the conglomeratic sedimentary unit (Tsc) (Plate 1), and the number and diameter of the blocks generally diminish to the south. The source area was most likely in the footwall of the major northwest-striking fault south of Mount Tobin, where only pre-Tertiary rocks are now exposed (Fig. 3). Although these blocks locally extend 5 km south of this fault in the southeast part of Plate 1, the breccia unit does not apparently extend farther south, and thick sections of breccia are confined to within ~2 km of the fault. In the southeast part of the map area, in most places the Miocene lacustrine and fluvial sedimentary unit directly overlies the ca. 33 Ma effusive andesite sequence. Where the top of the sedimentary and breccia units is present, it is overlain by the Miocene lacustrine and fluvial sequence. These observations suggest that most Fish Creek Mountains Tuff exposures in the Tobin Range are not in the site of primary deposition, but rather are secondary syntectonic landslide breccia deposits associated with normal faulting. An alternative explanation is that the Fish Creek Mountains Tuff megablocks are actually part of a debris flow derived from the Fish Creek Mountains caldera that flowed westward down the axis of the paleovalley, similar to models proposed for tuff megabreccia blocks in northeastern Nevada (Henry, 2008). Some of the Tobin Range megabreccias lack any matrix and contain 10 cm to 1 m Fish Creek Mountains Tuff blocks with rotated fiammé that are plastically deformed and sutured to one another, as if the blocks were emplaced shortly after eruption, when the tuff was hot and still plastic (Fig. 8C). On the geologic map (Plate 1), areas dominated by rhyolite landslide breccia (Trx) are distinguished from those areas dominated by clastic sedimentary rock (Tsc) by a generalized and approximately located contact.

Younger Rhyolite Ignimbrite (Tri)

Overlying the sedimentary rock and landslide breccia unit is an isolated exposure of moderately welded, crystal-rich, lithic-poor rhyolite ignimbrite with vapor phase mineralization near the top of the unit. Its age and correlation remain uncertain. A single geochemical sample contains 77 wt% silica as well as elevated Th (30 ppm) and Rb (300 ppm).

Lacustrine Sedimentary Rocks (Ts) and Younger Basalt (Tyb)

Stratigraphically above the younger rhyolite ignimbrite and landslide breccia is a thick sequence of lacustrine and fluvial sedimentary rocks consisting of ash- and pumice-rich tuffaceous siltstone, mudstone, and sandstone with a few thin beds of buff colored sandy limestone. Deffeyes (1959) provided a detailed study of these rocks in Jersey Valley and estimated a thickness of >820 m for this unit. Based on Barstovian and early Calendonian (middle and early-late Miocene) mammalian fossils within the unit, an age of 15–4 Ma has been estimated (Deffeyes, 1959). Interbedded in the lower to middle part of the unit is a series of basaltic lava flows containing 20 vol% plagioclase and 7–10 vol% olivine. A whole-rock 40Ar/39Ar plateau age of 14.10 ± 0.12 Ma was obtained for a basaltic lava. Additional basalt exposures are southwest of Needle Peak and consist of basaltic lava flows and dikes that overlie, bury, and perhaps intrude the major northwest-striking system of normal faults (Plate 1). The basalt is poorly exposed here, but lithologically resembles the lavas farther southeast. These basalt lavas dip gently and overlie in angular unconformity pre-Tertiary rocks on the north and 25° east-dipping Caetano Tuff and overlying early Oligocene andesitic lava flows on the south.

Quaternary Deposits (Q)

Quaternary units consist of alluvium, landslide deposits, and gravel deposits. Some gravel deposits consist of clasts primarily derived from the Pumpernickel and Havallah Formations and are localized near the base of the major northwest-southeast–striking normal fault. Other gravel deposits are exposed near the Quaternary Tobin Range fault and a second fault east of Needle Peak.


Normal faults in the Golconda Canyon region of the Tobin Range principally strike north-south and dip steeply to moderately to the west. Predominantly westward-dipping faults have produced eastward tilting, thus resulting in eastward-dipping Tertiary rocks. Three main groups of faults were identified and are described below. Figure 10 illustrates the sequence of faulting along sections A-A' and B-B' of Plate 1.

Early Oligocene Faults

Throughout the Tertiary paleovalley in Golconda Canyon, numerous small faults with dips ranging from ~35° to 80° to the west and northwest have been identified (Plate 1 and Fig. 7). These faults generally strike north or northeast, though occasionally strike east-west, and are truncated by a major northwest-striking fault. These faults are better exposed, more numerous, and have greater offset south of Golconda Canyon, where Burke (1973) documented that north-striking, moderately west-dipping faults offset the older Basalt and overlying Caetano Tuff as much as 500 m. As illustrated in Figure 3, these early faults are intruded by subvolcanic andesite intrusions that are undated, but appear to be cogenetic with the early Oligocene andesitic sequence (Burke, 1973; Tatlock et al., 1977). The andesitic lava flows and breccias dip 25°–40°E, and both bury and are not offset by some of these faults. The underlying and faulted Caetano Tuff generally dips slightly (~5°) more steeply to the east at angles of 35°–45° (Fig. 3). These relationships demonstrate that the faults are earliest Oligocene in age.

The timing of this set of faults in Golconda Canyon is best determined by relationships seen in the andesite sequence in the western part of the study area (Plate 1 and Fig. 11). The andesitic rocks dip from ~15° to 45° to the east. On the north side of Golconda Canyon the Caetano Tuff dips moderately to the east-northeast (~25°–35°) and is overlain by a series of andesitic ignimbrites and lava flows with slightly shallower dips that decrease upsection (~15°–25°). One particular fault offsets the basal ignimbrites and overlying pyroxene andesite lavas, but not the upper hornblende-bearing andesitic lava flows, and hence is buried by the latter. The andesitic ignimbrites and overlying pyroxene andesite lavas are intercalated with landslide breccias in the hanging wall of this fault (Fig. 11). These relationships indicate a period of faulting synchronous with and extending a short time after the ignimbrite volcanism, dated here as 33.03 Ma, but prior to eruption of the hornblende-bearing lava flows. In this area, this early stage of faulting resulted in ~5° to locally 10° of eastward tilting of the andesite ignimbrites and older rocks.

We conclude that the earliest Oligocene normal faulting was widespread but small in magnitude in the western part of the study area (Golconda Canyon) and farther south in the Tobin Range, and resulted in 5°–10° of westward tiling. In the southeast part of the study area, the Caetano Tuff and older rocks are sparsely exposed, so this age of faults is not documented.

Late Oligocene(?) to Middle Miocene Faults

The largest magnitudes of tilting and normal fault displacements are middle Miocene to possibly as old as late Oligocene(?). A major northwest-striking normal fault crosses the Tobin Range and both offsets and locally forms the northern boundary of the Tertiary paleovalley. This fault strikes ~N30°W for ~10 km west of the range crest, but bifurcates to the east (Fig. 3 and Plate 1). One strand bends to an east-west strike for 5 km to its termination on a Quaternary fault on the eastern front of the Tobin Range near Needle Peak. The second strand is poorly exposed, but extends in a S10°–20°E direction ~6 km as two fault traces to where it is buried beneath Miocene sedimentary rocks. The southern fault dips 50°W, offsets the lower part of the Miocene sedimentary section more than 500 m, and has been intruded after movement by a Miocene(?) dacite dike (Tdi; Plate 1).

In the northwest part of the map area, the northwestward-striking fault is easily mapped due to a prominent topographic break and juxtaposed lithologies. The fault dips ~50°–55°SW and juxtaposes the Triassic China Mountain Formation against a fault sliver of biotite rhyolite ignimbrite (Trib) and early Oligocene andesitic conglomerate and sandstone (Plate 1 and Fig. 9). Palinspastic restorations suggest that this fault has accommodated locally more than 1.7 km of normal displacement (Figs. 10A, 10B). Movement on this fault has apparently also localized much of the Tertiary sedimentary rock and landslide breccia unit in the hanging wall. The age of this fault system is constrained by the onset of sedimentary rock and landslide breccia deposition after the deposition of the Fish Creek Mountains Tuff (24.72 Ma) and prior to the eruption of the younger Tertiary basaltic lavas (14.10 Ma) and dikes, which respectively overlie and intrude the eastern end of the fault (Fig. 12). Based on the observation that the fault system offsets the lower part of the middle(?) Miocene lacustrine sedimentary sequence that apparently underlies the 14.10 Ma basaltic lavas dated here, the fault system was likely active in the early-middle Miocene (ca. 15 Ma). Andesitic sedimentary rocks in Bushee Canyon on the west side of the range dip 15°–30°E and give an estimate of tilting associated with this generation of faults (Plate 1 and Fig. 9A).

Several other north-south and north-northwest–striking faults may be similar in age to the northwest-striking fault described above. These faults are poorly exposed on the south side of the field area in the center of the range, where in places they contain landslide megabreccias in their hanging wall and in other places contain basal, limestone-bearing lacustrine sedimentary rocks underlying the basaltic lavas. One northwest-striking fault that cuts the base of the lacustrine sequence was intruded by a dacite dike.

Middle and Late(?) Miocene Faults

Most faults middle Miocene or younger in age strike north-south, dip moderately to the west, and offset the northwestward-striking middle Miocene faults (Plate 1) as well as the 14.10 Ma basalts that dip ~25°–30°E in the eastern part of the study area, as illustrated in cross section (Plate 1 and Fig. 10). The dips of the basalts are a result of movement on these faults as well as active faults along the east and west sides of Jersey Valley. The dips of these basalts are similar to the magnitude of the dips of early Oligocene volcanic rocks to the west, near the mouth of Golconda Canyon, that were affected by the older early Oligocene phase of faulting and the middle Miocene and older faults. This comparison suggests that this younger phase of normal faulting did not affect the western area as significantly as the area to the east, or that the western area of middle Miocene faulting migrated eastward over a short period of time to form slightly younger faults to the east. Regardless, post–14.10 Ma faults on the eastern part of the Tobin Range map area produced stratal tilts equivalent to and locally greater than all the post-Eocene faults on the western part of the range. Dips on the Caetano Tuff average ~25° eastward, compared to ~30° eastward dip of the middle and late Miocene sedimentary rocks as illustrated in Figure 13.

On the eastern side of the Tobin Range several steeply eastward-dipping faults appear to be middle Miocene in age. One particular fault that forms a contact between the Caetano Tuff and the andesite lava sequence has a northeastward strike and dips steeply to the east (Plate 1 and Fig. 12). This fault postdates and offsets the major late Oligocene–middle Miocene northwest-striking fault and has an orientation and displacement similar to the fault east of Needle Peak that bounds the eastern limit of the Tobin Range (see following). This fault is significant because displacement along east-dipping faults would cause westward tilt and diminish the magnitude of the eastward dip of older Tertiary rocks.

Quaternary Faults

A subset of the younger faults described above comprises faults that have evidence of Pleistocene or Holocene movement. The youngest faults include those bounding either side of the Tobin Range. The western limit of the Tobin Range is bounded by a large west-dipping normal fault that ruptured during the 1915 Pleasant Valley earthquake. The Pleasant Valley earthquake is the most northern surface-rupturing event in the Central Nevada seismic zone to have occurred during historical time. A detailed study by Wallace (1984a) documents the surface rupturing of this event. Average vertical displacement along the scarps was 2 m with a maximum displacement of 5.8 m occurring along the most southern scarp, near the mouth of Golconda Canyon (Wallace, 1984a). Lateral displacement was minimal, but when present was right lateral. Wallace (1984a) reported a maximum lateral displacement of 2 m just north of Golconda Canyon. Recurrence intervals for the Pleasant Valley fault are estimated between ~5 k.y. to <12 k.y. (Wallace, 1984b; Bonilla et al., 1984).

The eastern limit of the Tobin Range east of Needle Peak in the northeast part of the map area, and continuing north along the margin of Buffalo Valley, is bounded by a large normal fault dipping to the east. A degraded scarp in alluvium attests to Quaternary movement in Buffalo Valley. The displacement on this fault dies out as it extends south into Jersey Valley, where there are several scarps in Quaternary alluvium (Plate 1), and displacement appears to be transferred east across the valley to the normal fault bounding the western margin of the Fish Creek Mountains.

Normal Faulting and Tilting West of the Tobin Range

We did a reconnaissance examination of Oligocene ignimbrite exposures in an east-west transect from the Tobin Range to the West Humboldt Range, north of Lovelock (Fig. 14). Isolated exposures of late Eocene or Oligocene ignimbrites (of uncertain correlation) overlie Mesozoic rocks in this zone, and all are significantly tilted and associated with steeply to gently dipping normal faults that mainly dip to the west (cf. Johnson and Barton, 2000). Ignimbrite dips range from 20°E to 80°E, and most average 35°E to 40°E (see also data for the Granite Mountain and Sou Hills area; Fosdick and Colgan, 2008). Overlying Miocene(?) tuffaceous sedimentary rocks dip as much as 40°E in the Sou Hills flanking the southern Tobin Range, but elsewhere dip <20°E. Basaltic lavas are widespread and were dated by Fosdick and Colgan (2008) in the East Range and Sou Hills as 17–13 Ma; they are potentially similar in age to the 14.10 Ma basalts of the Tobin Range. These basalts overlie the sedimentary deposits or ignimbrites and display shallow eastward dips ranging from ~2°E to 15°E with an average of 10°E. On the basis of the dips of these rocks, we conclude that extensional normal faulting has decreased with time, but appears to have extended Oligocene and older units ~50%–100% (see methods below) between the Tobin Range and Lovelock. This conclusion is similar to the ~60% extension estimate by Fosdick and Colgan (2008) for the East Range and Sou Hills part of this larger area.


Timing and Amount of Extension

Extension in the Basin and Range Province has occurred by means of normal faults, and the youngest of these faults have produced the characteristic topography of horsts and grabens (Stewart, 1978; Spencer and Reynolds, 1986; Dickinson, 1991). Block rotation of rocks accompanies normal faulting and results in strata rotating the direction opposite from fault dip. A single generation of normal faults dipping steeply at 50°–60° and with a spacing of 10–15 km, such as the modern Basin and Range of Nevada, can account for 10%–20% total extension and 5°–15° block rotation (Stewart, 1978; Lerch et al., 2004; Colgan et al., 2006a). Areas that have undergone greater amounts of extension, in many places exceeding 100%, have undergone more complex normal faulting in which faults are either more closely spaced and/or movement has occurred on two or more successive generations of normal faults (Anderson, 1971; Proffett, 1977). Where each generation of normal faults dips in the same direction, successively crosscutting younger faults cause inactive older normal faults to rotate to shallower dips. The result is a highly extended terrane with shallowly dipping normal faults and steeply dipping rock units. Areas of high extension have also been described together with the presence of low-angle detachment faults (cf. Davis and Lister, 1988); however, detachment faults have not been identified in most areas that have undergone large amounts of extension in volcanic terranes of western Nevada, such as the Yerington district and the El Dorado Mountains (Anderson, 1971; Proffett, 1977).

For a single generation of normal faults with the same dip direction, the amount of extension can be calculated from the average dip of the rock units and an estimate of the initial dip of the normal faults by using the following equation from Thompson (1960):  
where e is the amount of extension in percent, θ is the original dip of normal faults, and ϕ is the dip of rock units that were flat prior to faulting.

The Tertiary rock units in the Golconda Canyon area have dips that range between 15° and 40° to the east. The average dip in the area is ~25° (Fig. 13), and this amount of dip is relatively consistent throughout the range. If we assume that all of the normal faults initiated with a dip of 60° and use an average dip of 25° for the rock units, east-west crustal extension in the Tobin Range is calculated to be 51%. This calculation provides a rough estimate but does not adequately accommodate multiple generations of faults or situations where some faults dip in the opposite direction to the main fault sets and therefore reduce dip of rock units. In the former case, the Thompson method overestimates the amount of extension, whereas in the latter case the Thompson method underestimates the amount of extension.

A more accurate estimation of the amount of extension in the Tobin Range can be obtained from palinspastic restorations of cross sections drawn parallel to the slip direction along the faults. Cross-section restorations provide extension estimates between 25% and 46% (Fig. 10). Because of factors such as erosion and faults for which the amount of displacement cannot be determined accurately, the palinspastic estimations provide a minimum percentage of extension. For example, the restored cross-sections A-A' and B-B' do not include faults of earliest Oligocene age that are significant in the Tobin Range south of these sections, and section B-B' has been affected by east-dipping Quaternary normal faults along the east flank of the study area that have reduced the eastward stratal dips. We therefore estimate that the southern Tobin Range has undergone ~50%, or possibly slightly more, extension.

The timing of extension and magmatism in the Tobin Range is summarized in Figure 15. The earliest recognized normal faulting and minor extension (up to 5°–10°E stratal tilts) occurred in a short period of time synchronous with eruption of the earliest andesitic rocks at 33 Ma, and as followed by a period of little or no faulting from ca. 33 Ma to 24.5 Ma. The largest amount of extension, associated with the landslide breccias, is bracketed between the 24.7 Ma age of the Fish Creek Mountains Tuff, and eruption of basaltic lavas at 14.10 Ma. Although the timing of extension cannot be bracketed more closely than this 10 m.y. interval, we infer that the faulting likely took place toward the end of this period and extended into the late Miocene, because the ca. 15–4 Ma Miocene sequence of lacustrine and fluvial sedimentary deposits overlying the breccias is faulted and tilted up to 40°E. The currently active normal faults may have extensional rates similar to or slightly lower than those of the middle Miocene. We infer that ~80% of the extension occurred in middle Miocene time (ca. 20–10 Ma) with a lesser amount of faulting post–10 Ma and ca. 33 Ma.

Regional Extension Comparisons

Estimates of extension in the Tobin Range agree with a pattern of decreasing amounts of extension to the northwest, in northwest Nevada (Fig. 2). Regional palinspastic restorations by Smith et al. (1991) suggest total amounts of crustal extension between 50% and 100% in central Nevada. Numerous studies indicate that early (pre–10 Ma) extension is spatially heterogeneous, with areas of extreme extension separated by zones of little-extended crust (cf. Gans and Miller, 1983; Gans, 1987; Smith, 1992; Dilles and Gans, 1995). John (1995) and Hudson et al. (2000) estimated >100% extension ca. 25 Ma for the Job Canyon caldera complex in the nearby southern Stillwater Range, and Colgan et al. (2008) estimated 100% middle Miocene extension for the Caetano caldera east of the Tobin Range. Using Thompson's (1960) method for estimating extension, our reconnaissance west from the Tobin Range to the West Humboldt Range suggests an average of 13% extension postdating Miocene basalt lavas, and 50%–75% extension between Oligocene ignimbrite deposition and Miocene basalt eruption, based on tilting of basalts ~10° and older ignimbrites an additional 40°. These estimates agree with the 50% extension documented by Fosdick and Colgan (2008) for the Granite Mountain area (Fig. 14). To the northwest, workers have found extension of ~10%–20% in the Black Rock and Pine Forest Ranges (Lerch et al., 2004; Colgan et al., 2006a, 2006b). Likewise, central Nevada and northwest Nevada display a difference in timing in normal fault inception (Fig. 11). In the Tobin Range normal faulting initiated with minor extension during the early Oligocene, followed by at least two generations of faulting beginning in the Miocene. The minor extension associated with intermediate magmatism of early Oligocene age is not documented elsewhere outside the Tobin Range in northwestern Nevada, but has been noted in the Tuscarora and Carlin districts of north-central Nevada (Fig. 2; Castor et al., 2003; Ressel and Henry, 2006). As noted by Henry (2008, cf. Fig. 1), such faults were minor enough to not disrupt the kilometer-deep paleovalleys and associated topography of Eocene and Oligocene age that characterize much of northern Nevada. The Tobin Range observations are consistent with work by Smith (1992) and Hudson et al. (2000), who identified at least two generations of normal faulting, beginning in the Oligocene in the Toiyabe and Stillwater Ranges. The timing of the main pulse of middle Miocene extension between ca. 20 and 10 Ma in the Tobin Range is similar to the timing of the ~100% extension of the Caetano caldera of the Shoshone Range between 16 Ma and 10–12 Ma (Colgan et al., 2008) and to the timing of ~50% extension in the Granite Mountain area of the East Range between 15–17 Ma and 10 Ma (Fosdick and Colgan, 2008). In contrast, in northwesternmost Nevada, Colgan et al. (2006a, 2006b) identified a single generation of normal faulting that began some time after the late Miocene.

Regional Migration of Extension

Normal faulting and extension in Nevada have migrated with time, such that the older faulting episodes are found in northeastern Nevada, and with time the zone of maximum faulting and extension rates has generally migrated south, west, and east (Seedorff, 1991). Late Eocene (35–40 Ma) extension in northeastern Nevada is closely associated with the initiation of dacitic magmatism (cf. Gans et al., 1989; Feeley and Grunder, 1991) and with Carlin and epithermal-style gold mineralization (Ressel and Henry, 2006; Castor et al., 2003). Oligocene extensional faulting documented for the Toiyabe Range, Still-water Range, and Atlanta district are associated in time with caldera complexes that are the source of large silicic ignimbrite sheets. In the Walker Lane belt between Las Vegas and Reno, extension, normal faulting, and tilting began locally in the late Oligocene, whereas widespread but patchy zones of extension associated with andesitic magmatism began ca. 15 Ma (cf. Dilles and Gans, 1995; Surpless et al., 2002). Data reported by Colgan et al. (2008), Fosdick and Colgan (2008), and this study indicate that this middle Miocene extension also extended northward into north-central Nevada. In the past 10–15 m.y., faulting has moved westward to the edge of the modern Sierra Nevada block and has propagated north-westward (cf. Henry et al., 2004; Carmichael et al., 2006; Colgan et al., 2006b). At present, the most active faulting in Nevada is concentrated along the Sierra Nevada front, in the Walker Lane belt, and in the Central Nevada seismic zone that extends northward from the Walker Lane through the western margin of the Tobin Range.

Significance of Mid-Tertiary Paleovalleys

The Eocene paleovalley documented herein for the southern Tobin Range is similar in age and orientation to several others documented in the northern Sierra Nevada, California (Lindgren, 1911), in west-central Nevada (Proffett and Proffett, 1976; Henry et al., 2003), and in northeastern Nevada (Henry, 2008). The Tobin Range paleovalley provides evidence that the northeastern Nevada and western Nevada–Sierra Nevada paleovalleys were linked as part of a system of river drainages (cf. Henry, 2008).

In all these areas, the paleovalleys with 1–2 km of topographic relief were developed via incision of river systems into a stable plateau prior to onset of extensional normal faulting (Henry, 2008). In the Tobin Range of northeastern Nevada, and the Yerington district of western Nevada, the paleovalleys persisted after the onset of local or small-magnitude extension dated ca. 33 Ma, and ca. 26–23 Ma, respectively, in these two localities (this study; Henry, 2008; Dilles and Gans, 1995). Middle Miocene rapid and high-magnitude extension disrupted and terminated the paleovalley system in the Tobin Range and the Yerington district (Proffett, 1977; Dilles and Gans, 1995).

Relationship to Gold Mineralization

Nevada is a leading producer of gold from Carlin-type deposits associated with late Eocene magmatism (cf. Ressel and Henry, 2006) and from epithermal veins and disseminated deposits associated locally with Oligocene caldera complexes (cf. Round Mountain) and more generally with middle Miocene andesitic volcanism of the Walker Lane belt. Magmatism and normal faulting are common ingredients.

Although the only significant mineralization in the Tobin Range is the Mount Tobin mercury mine (Johnson, 1977), this mineralization is along the Miocene fault system that produced the landslide megabreccias. More significantly, the data from the southern Tobin Range further document association of normal faulting with magmatism, which began here in the early Oligocene and extended into the middle Miocene (Fig. 15). Furthermore, the Tobin Range and reconnaissance data in the ranges to the west indicate these areas have in general undergone 50% and locally as much as 100% extension. These data fill in a gap and suggest that zones of moderate to high extension extend in a patchy manner from northeast Nevada southwest to the Walker Lane, in contrast to the little extended zones of far northwestern Nevada documented by Colgan et al. (2006a, 2006b).


Cenozoic extensional normal faulting has occurred since the early Oligocene in the Golconda Canyon area. The first phase of normal faulting was small in magnitude (5°–10° stratal tilting) and developed coeval with the eruption of the andesite ignimbrites and lavas (ca. 33 Ma) and the intrusion of andesitic dikes. This early Oligocene extension is consistent with the regional pattern of westward-younging normal fault inception ca. 35–40 Ma in east-central and northwest Nevada (Fig. 2). During this phase, faults were principally localized in the western part of the range south of Golconda Canyon (Fig. 3). The second phase of faulting resulted in the development of a moderately west-dipping system of northwest-striking faults that was active from ca. 20 Ma(?) to 14 Ma in the middle Miocene and possibly the late Oligocene, but predated eruption of 14.1 Ma basaltic lavas. This fault system was responsible for the localization of syntectonic clastic sedimentary rocks and landslide megabreccias with rhyolite ignimbrite clasts, which accumulated in the hanging walls of the faults. A series of faults younger than basalts of middle Miocene age (14.1 Ma) postdates the northwest-trending fault system but is possibly temporally continuous with the earlier faults. These younger faults are exposed in and possibly largely confined to the eastern part of the range. These two sets of Miocene faults produced ~20°–25° stratal tilts and ~70%–80% of the extension in the Tobin Range over the period ca. 20 Ma(?) to ca. 10 Ma. The timing of the Miocene extension in the Tobin Range is similar to the recently documented extension to the west and east in the nearby East Range and Shoshone Range and to the broad zone of extension in the Walker Lane of western Nevada (Fig. 2). The youngest set of normal faults (younger than ca. 10 Ma) is active today in the form of moderately to steeply dipping range front faults that bound either side of the Tobin Range.

Total crustal extension for the Tobin Range is estimated to be ~50% or slightly more. This estimate of extension is consistent with other estimates for central Nevada and areas farther east, as opposed to areas to the northwest, where extension diminishes to 10%–20%.

John Huard and Robert Duncan performed the Ar-Ar age determinations reported herein, and Chris Henry kindly obtained the Ar-Ar age for the Caetano Tuff in Golconda Canyon. We thank Chris Henry, David John, Joe Colgan, and Tom Moore for sharing their ideas, data, and interests in our work and their parallel studies. Robert Yeats and Andrew Meigs helped guide this study, and Joe Becker supplied field assistance. This work was partly funded by an Edmap grant from the U.S. Geological Survey, and represents a collaboration between Oregon State University, the U.S. Geological Survey, and the Nevada Bureau of Mines and Geology. We thank Joe Colgan, Ben Surpless, and Alan Wallace for their thorough and helpful reviews of this manuscript.

If you are viewing the PDF of this paper or reading it offline, please see the full-text article on www.gsajournals.org to view Supplemental Tables 1, 2, 3, 4, 5, 6, and 7. Or, you may view each table individually by visiting: http://dx.doi.org/10.1130/GES00137.S1 (Supplemental Table 1); http://dx.doi.org/10.1130/GES00137.S2 (Supplemental Table 2); http://dx.doi.org/10.1130/GES00137.S3 (Supplemental Table 3); http://dx.doi.org/10.1130/GES00137.S4 (Supplemental Table 4); http://dx.doi.org/10.1130/GES00137.S5 (Supplemental Table 5); http://dx.doi.org/10.1130/GES00137.S6 (Supplemental Table 6); http://dx.doi.org/10.1130/GES00137.S7 (Supplemental Table 7).