Geologic mapping, supported by 40Ar/39Ar and U-Pb geochronology and geochemistry of sedimentary and volcanic rocks, reveals the details of the Cenozoic depositional and tectonic history of the eastern Piñon Range and central Huntington Valley in the north-central Basin and Range Province, Nevada (USA). Cretaceous to Miocene supracrustal successions were studied in detail in order to compare the geologic evolution of the upper crust near the Ruby Mountains–East Humboldt Range (RMEH) metamorphic core complex (MCC) with the magmatic, metamorphic, and deformational history of the deep crust in the developing MCC. During the well-documented Late Cretaceous–Oligocene history of partial melting and infrastructure development within the RMEH, surface deposits in Huntington Valley reflect general tectonic quiescence, with evidence for the development of the shallow Elko Basin, minor extension, and eruption of southward-younging ignimbrite flare-up volcanism. Thin, discontinuous successions of Cretaceous–early Cenozoic sedimentary strata were locally blanketed by rhyodacite ignimbrites, domes, and subvolcanic intrusions of the Robinson Mountain volcanic field between 38.5 and 36.8 Ma. This magmatic event represents the first local expression of Cenozoic volcanism linked to the ignimbrite flare-up, and its onset occurred slightly after a renewal of partial melting in the RMEH beginning ca. 42 Ma. The volcanic section was subsequently tilted ∼10°–15° west before ca. 33.9 ± 0.4 Ma. Although melting continued at depth in the RMEH until after 30 Ma, there was no eruption of volcanic rocks after Robinson Mountain volcanism. An additional ≥10°–15° of westward tilting occurred between 31.1 ± 0.3 Ma and ca. 24.4 Ma, as bracketed by the 31 Ma tuff of Hackwood Ranch (which was probably erupted from a distant caldera) and an angular unconformity beneath the overlying Miocene Humboldt Formation. Neither of these tilting events and unconformities appears to represent significant (>∼1 km each) extension, but they could be surface expressions of magmatism, metamorphism, and crustal flow at depth. The Humboldt Formation includes >2000 m of sediment deposited mostly between ca. 16 and 12 Ma, with deposition lasting until at least ca. 8.2 Ma. Humboldt Formation sediments thicken eastward, toward the west-dipping fault that bounds the RMEH, and are interpreted as a basin that developed in the hanging wall of this fault system. Motion on this normal fault system led to the exhumation of metamorphic and igneous rocks of the core complex ∼10 m.y. after the documented cessation of partial melting, high-temperature metamorphism, and intrusion of granitoids into the deep crust ca. 29 Ma. Metamorphic clasts and a detrital zircon signature thought to represent RMEH provenance are first detected in 14.2 Ma or younger sediments.
The northern Basin and Range Province (BRP) of western North America (Fig. 1A) has had a protracted history of Cenozoic east-west–directed extension that also involved the uplift of anomalous exposures of deeper crustal rocks in metamorphic core complexes (MCCs) (e.g., Coney, 1980; Armstrong, 1982; Coney and Harms, 1984; Dickinson, 2006, 2013; Colgan, 2013). Despite decades of study, there is still much disagreement on the exact timing, nature, and geodynamic setting for the uplift of metamorphic rocks in MCCs (e.g., discussions in Wells et al., 2012; Miller et al., 2012).
Quantitative geobarometry of metamorphic rocks in northern BRP core complexes yields data that require uplift of some of these rocks from at least 35 km depth, perhaps in more than one stage of extension, beginning in the Late Cretaceous to early Cenozoic (e.g., Hodges et al., 1992; McGrew and Snee, 1994; Lewis et al., 1999; McGrew et al., 2000; Wells and Hoisch, 2008; Cooper et al., 2010; Wells et al., 2012; Hallett and Spear, 2014, 2015; Willis, 2014). These data have implications for crustal thicknesses at the end of crustal shortening in the Mesozoic and whether the crust had become thick enough to collapse gravitationally prior to Basin and Range extension, as proposed by Vandervoort and Schmitt (1990), Hodges and Walker (1992), Camilleri and Chamberlain (1997), Sonder and Jones (1999), DeCelles and Coogan (2006), Wells and Hoisch (2008), and Wells et al. (2012). The major structures proposed by these authors to have operated during the Cretaceous to Paleogene extensional events have not been clearly identified as surface-breaking features except in one place (Camilleri, 1996; Camilleri et al., 1997). Maps of rock units and structures beneath the reconstructed Cenozoic unconformity have also not found clear evidence for these structures (e.g., Gans and Miller, 1983; Van Buer et al., 2009; Konstantinou et al., 2012; Long, 2012). Some argue that, near the Ruby Mountains–East Humboldt Range (RMEH), motion on the west-dipping MCC detachment fault (Fig. 1) began at or before the Eocene or Oligocene (e.g., Dallmeyer et al., 1986; Dokka et al., 1986; McGrew and Snee, 1994; Howard, 2003), and the accumulation of Paleogene or older strata next to the Ruby Mountains has been used as evidence for significant pre-Miocene normal faulting (e.g., Solomon et al., 1979; Vandervoort and Schmitt, 1990; Mueller and Snoke, 1993; Brooks et al., 1995; Satarugsa and Johnson, 2000; Rahl et al., 2002). However, Henry (2008), Colgan et al. (2010), and Henry et al. (2012) provided evidence that RMEH MCC uplift happened almost entirely in the Miocene with negligible faulting prior to this time (see discussion in Henry et al., 2011), a discovery compatible with the Miocene uplift histories of other core complexes of the northern BRP (e.g., Miller et al., 1999; Konstantinou et al., 2012, 2013; Ruksznis, 2015).
This study directly addresses these problems and questions by documenting in detail the evolution of the upper crust across a broad region flanking the RMEH (Figs. 1B, 2, and 3) to investigate which events are recorded by its rich and long-lived Late Cretaceous and Cenozoic sedimentary and volcanic history, and how this evolution compares to the coeval history of the deeper crust represented by the igneous and metamorphic rocks of the RMEH. Our contribution is based on 1:24,000 scale geologic mapping of sedimentary and volcanic rocks present within central Huntington Valley and the eastern Piñon Range (Fig. 2) west of the southern and central RMEH. This region is underlain by a remarkably complete and well-exposed Cenozoic succession (Fig. 1B). Previous work (Smith and Ketner, 1976, 1978; Horton et al., 2004) reported thick accumulations of Late Cretaceous–Quaternary rocks, some of which have been previously interpreted to represent pre-Miocene synextensional sedimentation and volcanism associated with evolution and exhumation of the RMEH (Satarugsa and Johnson, 2000; Haynes, 2003; Howard, 2003; discussion in Henry et al., 2011). Our study of this same stratigraphy incorporates extensive detrital zircon U-Pb geochronology, 40Ar/39Ar ages of sanidine from tuffs, and U-Pb zircon geochronology and whole-rock geochemistry of volcanic rocks in order to better constrain depositional ages and improve lithologic correlations within this succession. Our new data allow us to more carefully assess the supracrustal history and compare it to the magmatic, metamorphic, and deformational history detailed for the RMEH MCC. The data also significantly improve our understanding of regionally important Cenozoic time-stratigraphic sections, some of which have been targeted for Cenozoic paleoclimate and paleoelevation studies (e.g., Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Feng et al., 2013; Mulch et al., 2015).
The RMEH exposes Precambrian basement, Neoproterozoic–Paleozoic metamorphosed shelf margin sediments, and Mesozoic–Cenozoic plutonic rocks (e.g., Snoke, 1980; Stewart, 1980; Wright and Snoke, 1993; MacCready et al., 1997; McGrew et al., 2000; Premo et al., 2008, 2014; Colgan et al., 2010; Lund et al., 2010; McGrew and Snoke, 2010, 2015; Howard et al., 2011). These rocks were exhumed by Cenozoic normal slip on a (now) shallowly west-dipping (∼20°) brittle fault system mapped along the western side of the RMEH that cuts a subparallel mylonite shear zone (e.g., Mueller and Snoke, 1993; Sullivan and Snoke, 2007; Colgan et al., 2010). Continued uplift of the range occurred along steeply west- and east-dipping normal faults bounding the west and east sides of the range (Colgan et al., 2010; Fig. 1B).
Most of the pre-Cenozoic strata exposed in the Carlin-Piñon Range and surrounding region (Fig. 1) consist of ocean basin and continental slope sedimentary and volcanic rocks of the Roberts Mountains allochthon, emplaced over shelfal rocks during the earliest Mississippian Antler orogeny (Burchfiel and Davis, 1975; Johnson and Pendergast, 1981; Turner et al., 1989; Trexler and Nitchman, 1990; Burchfiel et al., 1992; Johnson and Visconti, 1992; Miller et al., 1992; Trexler et al., 2004). The study area (Fig. 1B) contains foreland basin strata deposited east of the Antler thrust front, including the Chainman Shale and Diamond Peak Formation sandstone and conglomerate (Smith and Ketner, 1978; Trexler and Nitchman, 1990; Longo et al., 2002; Colgan and Henry, 2009). Pennsylvanian–Permian terrestrial and shallow-marine carbonates of the Antler overlap sequence were subsequently deposited across the eroded Roberts Mountains allochthon, foreland basin succession, and lower plate autochthonous rocks (Smith and Ketner, 1975, 1978; Stewart, 1980; Johnson and Pendergast, 1981; Johnson and Visconti, 1992; Ketner, 1998; Cook and Corboy, 2004; Colgan and Henry, 2009). Specific lithologies in the map area include quartzite and chert pebble conglomerates of the Mississippian–Pennsylvanian Diamond Peak Formation, shale and sandstone of the Mississippian Chainman Shale, silty, cliff-forming, cherty limestones of the Pennsylvanian Moleen Formation, and undivided platy siltstones and sandstones of Pennsylvanian–Permian age (Smith and Ketner, 1975, 1978; Stewart, 1980; Coats, 1987).
Regional pre-Cenozoic subcrop maps constructed across the northern Great Basin west of the Sevier fold and thrust belt indicate that mostly Pennsylvanian, Permian, and Triassic age strata were exposed at the surface prior to the eruption of Eocene–Oligocene volcanic rocks (Gans and Miller, 1983; Van Buer et al., 2009; Konstantinou et al., 2012; Long, 2012). Minor exposures of Cretaceous to Eocene strata occur beneath the volcanic section in isolated localities across eastern Nevada (e.g., Fouch et al., 1979; Vandervoort and Schmitt, 1990; Rahl et al., 2002; Druschke et al., 2009, 2011), but the Elko Basin (including the study area; Figs. 1 and 2) is one of the few places that has significant exposures of strata in this age range (Smith and Ketner, 1976; Solomon et al., 1979; Server and Solomon, 1983; Ketner and Alpha, 1992; Haynes, 2003; Henrici and Haynes, 2006; Henry, 2008; Colgan and Henry, 2009; Colgan et al., 2010; this study). The generally low temperatures of diagenesis for strata beneath the unconformity, based on organic maturation and conodont alteration indices (Gans et al., 1990; Crafford, 2005; Long, 2012) and low-temperature thermochronology (e.g., Colgan et al., 2010), indicate that they were never buried beneath substantial accumulations of sediment or tectonic overburden that might have been subsequently eroded away. However, one area, which includes the East Humboldt Range, Wood Hills, and Pequop Range, exposes strata exhibiting high metamorphic grades and yielding high geobarometric estimates (e.g., Camilleri, 1996; Camilleri and Chamberlain, 1997; McGrew et al., 2000; Hallett and Spear, 2014, 2015; Willis, 2014). It has been proposed that the Paleozoic rocks exposed in this area were buried during Jurassic–Cretaceous time beneath a thrust sheet, so this area would thus represent a proposed exception to the lack of tectonic burial elsewhere (Camilleri and Chamberlain, 1997; discussion in Long, 2012).
Volcanic activity began in the RMEH area of northeast Nevada ca. 42–40 Ma as part of the southward-sweeping volcanic front across Nevada (e.g., Armstrong, 1970; McKee et al., 1970, 1971; Armstrong and Ward, 1991; Best and Christiansen, 1991; Brooks et al., 1995; Humphreys, 1995; Ressel and Henry, 2006; Howard et al., 2011; Konstantinou et al., 2012). Voluminous volcanic and subvolcanic rocks in the study area (Fig. 2) record the arrival of the volcanic front ca. 38 Ma (Palmer et al., 1991; Gordee et al., 2000; Ressel and Henry, 2006; this study). The volcanic rocks are overlain by a thick succession of post-volcanic Miocene sedimentary rocks (Figs. 1–3) deposited in a fluviolacustrine setting with intercalated vitric tuffs derived from silicic eruptions in the Snake River Plain (Sharp, 1939; Perkins et al., 1998; Wallace et al., 2008; Colgan et al., 2010; Ellis et al., 2010; this study).
Geologic mapping of Cenozoic rocks in the central Huntington Valley and the eastern Piñon Range was carried out at 1:24,000 scale (Lund Snee and Miller, 2015), and a simplified geologic map is included here (Fig. 2). Photographs, photomicrographs, and detailed lithologic descriptions are also presented in Lund Snee (2013) and summarized in Table 1. Within the Eocene Robinson Mountain volcanic field, we differentiated igneous rock types and, where possible, mapped individual cooling units. In most cases we retained the geologic unit names for sedimentary rocks used in previous studies (Smith and Howard, 1977; Smith and Ketner, 1978), but we consolidated some units and eliminated the Smith and Ketner (1978) Indian Well Formation, much of which, based on this study, now forms part of the mostly Miocene Humboldt Formation.
The U-Pb and 40Ar/39Ar geochronology was carried out on 24 samples to better constrain depositional ages. The results of these analyses are summarized in Table 2, together with ages for parts of the section obtained from previous studies. Ion microprobe analyses were undertaken at the Stanford–U.S. Geological Survey Microanalysis Center to acquire trace element and U-Pb data from volcanic and subvolcanic units. Crystallization ages were determined from Tera-Wasserburg concordia intercept ages defined by the youngest zircon population (Fig. 4). Single crystal fusion 40Ar/39Ar results from alkali feldspar were obtained at the Stanford Noble Gas Laboratory, and weighted mean ages were used to calculate eruptive ages (Fig. 5). Inverse isochron ages for these samples generally agree with the weighted mean ages within 95% confidence (Supplemental File1). Errors for 40Ar/39Ar ages quoted in the text and Table 2 are 2σ total errors that facilitate comparison with U-Pb ages. These total errors take into account the systematic errors in flux monitor 40Ar*/40K (Phillips and Matchan, 2013) and 40K decay constants (Renne et al., 2010), in addition to the analytical errors. We used laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) methods to acquire detrital zircon age distributions at laboratories at the University of California, Santa Cruz and the University of Arizona LaserChron Center. Detrital zircon depositional ages were determined by taking the weighted mean of the youngest grains whose ages were indistinguishable at 95% confidence (Fig. 6). Detrital zircon age distributions are shown in Figure 7. Whole-rock geochemical analyses of 15 igneous samples were carried out at Macalester College (Minnesota) and Washington State University. A complete description of analytical details is included with data tables and supplementary plots of results in the Supplemental File.
CENOZOIC STRATIGRAPHY OF CENTRAL HUNTINGTON VALLEY AND THE EASTERN PIÑON RANGE, AND CONSTRAINTS UPON DEPOSITIONAL AGES
Late Cretaceous(?)–Eocene(?) Sedimentary Rocks
Conglomerate, sandstone, siltstone, and limestone are the oldest strata deposited above the Cenozoic unconformity developed across late Paleozoic strata. Pebble conglomerate and redbed sandstone occur at the base of the section (unit TKcs; Figs. 2 and 3A; Table 1). The redbeds are overlain by dense, brownish-purple, light orange, or white, massive to bedded limestone that commonly contains chert nodules (unit TKl; Fig. 3A). Although some (Smith and Ketner, 1976, 1978) differentiated the redbeds and overlying limestones into four map units, we follow Solomon et al. (1979), who grouped these strata into two units near the Elko Hills (Fig. 1B), TKcs and TKl (Fig. 2). TKcs and TKl have maximum thicknesses of ∼580 m and ∼600 m, respectively, and are stratigraphically below the Eocene Elko Formation (Fig. 2). Neither of these older units is present near Robinson Mountain, on the west side of the map area, where the younger Elko Formation directly overlies Mississippian- and Pennsylvanian-age rocks beneath the Cenozoic unconformity (Smith and Ketner, 1978; Figs. 1B and 2).
Conglomerates in the redbed succession (TKcs) have pebbles and cobbles of carbonate derived from underlying Paleozoic strata. Fine sandstones and siltstones are more common upward in the stratigraphic section, suggesting a transition from alluvial fan or fluvial to shallow lacustrine settings. A sample of pebbly sandstone (ELM11-PN16) collected from the redbeds (unit TKcs; Figs. 2 and 3A) yielded dominant U-Pb detrital zircon age peaks of 240 Ma, 380 Ma, 456 Ma, 629 Ma, 993 Ma, and 1.39 Ga, with minor maxima at 420 Ma, ca. 1.10 Ga, and 1.74 Ga (Fig. 7). The lack of Cenozoic zircons is striking compared to their abundance in younger strata (Fig. 7), and this supports Smith and Ketner’s (1976) decision to differentiate this unit from overlying strata such as the Elko Formation based on the lack of Cenozoic volcanic detritus in the redbeds. Although the lack of zircon younger than Triassic in this sample prevented us from calculating a maximum depositional age, we conclude that the redbeds and overlying limestones were most likely deposited before late Eocene (ca. 46 Ma or later) volcanic detritus was deposited in the area. Limited fossil evidence suggests that TKcs is probably not older than Cretaceous (Smith and Ketner, 1976). Based upon lithologic comparisons with other basal Cenozoic or Late Cretaceous clastic sedimentary rocks and carbonates described across the region, such as the Sheep Pass Formation (Fouch et al., 1979; Moore et al., 1983; Ketner and Alpha, 1992; Druschke et al., 2009, 2011), we provisionally assign a Late Cretaceous to Eocene (pre–46 Ma) age for deposition of both the redbeds (TKcs) and limestones (TKl).
Sample ELM11-PN16 has a detrital zircon signature (Fig. 7) most strongly resembling that of Mesozoic and late Paleozoic strata of northern Nevada and Utah. Detrital zircon age maxima exhibited by sample ELM11-PN16 are similar to those exhibited by the Chinle Formation, including peaks ca. 240 Ma, 456 Ma, 1.15–0.95 Ga, 1.5–1.35 Ga, and 1.8–1.6 Ga (Dickinson and Gehrels, 2008; Gehrels and Pecha, 2014; Figs. 7 and 8). The zircon signature of ELM11-PN16 thus likely represents recycling of zircons from the erosion of Triassic miogeoclinal strata that were exposed in the greater Elko Basin region at the end of the Mesozoic and beginning of the Cenozoic (Konstantinou et al., 2012; Long, 2012). Many of the observed age peaks in ELM11-PN16 are, in addition, shared with those in older Permian sedimentary rocks (e.g., samples PER-SS and DC-1 of Konstantinou et al., 2012) collected northeast of the study area (Figs. 1A and 8). The detrital zircon signature of Late Cretaceous(?)–Eocene(?) redbeds represented by sample ELM11-PN16 also appears to match the detrital zircon signature of the Pennsylvanian Spray Lakes Group and Permian Kindle Formation of British Columbia, particularly because these northern strata share a strong peak between 490 and 400 Ma (Gehrels and Pecha, 2014; Fig. 8). The lack of Mesozoic zircon in the Late Cretaceous(?)–Eocene(?) redbeds indicates no contribution from the Jurassic to Cretaceous Sierra Nevada batholith, which was undergoing uplift and erosion at this time (Van Buer et al., 2009), or from more proximal Mesozoic plutons (du Bray, 2007). This suggests that river systems draining Mesozoic granitic basement to the west flowed elsewhere and that local Mesozoic plutons, intruded at depth in the stratigraphic succession, were not yet sufficiently uplifted to be exposed at the surface.
A number of ages determined for strata within the overlying Elko Formation may have mistakenly been ascribed to the TKcs redbed conglomerates, which lack Cenozoic volcanic detritus (Horton et al., 2004; Mulch et al., 2015). These include a 46.1 ± 0.1 Ma U-Pb zircon age (sample 00-188GS in Table 2) obtained by Haynes et al. (2002) and Haynes (2003) on an air fall tuff from the Elko Formation in the Elko Hills (Fig. 1B). A ca. 43 Ma K-Ar age was also obtained from the same tuff (cf. Solomon et al., 1979; Haynes, 2003; Henrici and Haynes, 2006; sample BJS-1 in Table 2) as the 46 Ma zircon sample. Mulch et al. (2015, p. 322–324) mistakenly used these ages from the Elko Formation together with their new 40Ar/39Ar ages to extrapolate a 46.1–42.5 Ma depositional age range for portions of the underlying “cherty limestone” unit (our TKl). Horton et al. (2004, p. 868, 880, 882) cited an age range of 54.5–42.1 Ma for “limestone and limestone clast conglomerate,” “cherty limestone,” and “conglomerate, sandstone, siltstone” units. Both of the Horton et al. (2004) limestone units appear to be part of our mapped TKl, and their clastic unit is probably our TKcs. All of these units underlie the Elko Formation. In summary, the stratigraphic position of the limestones (TKl) and redbeds (TKcs) below the Elko Formation, together with their lack of volcanic material, suggests that they are older than 46 Ma, the oldest age determined on the Elko Formation, but how much older is unclear. Implications posed to stable isotope studies by this misclassification of ages for TKcs and TKl are discussed herein.
Eocene Elko Formation
The Eocene Elko Formation (unit Te) is stratigraphically above the Late Cretaceous(?)–Eocene(?) limestone (TKl) but directly overlies Paleozoic rocks across an angular unconformity in other parts of the mapped area (Figs. 2 and 3). The Elko Formation is a generally fining-upward succession of boulder to pebble conglomerate, sandstone, siltstone, limestone (commonly with chert nodules), laminated paper shale, and claystone with tuffaceous horizons and oil shale deposited within a generally fluvial-lacustrine environment (Smith and Ketner, 1976; Solomon et al., 1979; Moore et al., 1983; Server and Solomon, 1983; Ketner and Alpha, 1992; Haynes, 2003; Cline et al., 2005; Henrici and Haynes, 2006). In the mapped area, the Elko Formation is ∼180 m thick (Fig. 3B). Pebble conglomerate is exposed only immediately above the basal unconformity with Paleozoic rocks at one location (Figs. 2 and 3B), and paper shale and mudstone are interbedded with coarse clastic strata in the lower part of the section. Overall, the basal pebble conglomerate, sandstone, siltstone, and limestone fine upward to shale and interbedded tuff.
Sample ELKO-1 was selected from near the mapped base of the Elko Formation (Figs. 2 and 3B). It yielded a maximum depositional age of 45.92 ± 0.95 Ma, calculated from 20 of 100 detrital zircon grains (Fig. 6A; Table 2). Major age peaks occur at 156 Ma, 251 Ma, 429 Ma, 571 Ma, and 1.04 Ga and smaller peaks are at 1.44 Ga and 1.79 Ga (Fig. 7). Sandstone (sample ELKO-3) collected slightly higher in the Elko Formation (Figs. 2 and 3B) yielded a maximum depositional age of 45.00 ± 0.48 Ma on the basis of 32 of 37 analyses (Fig. 6B; Table 2). Additional minor age maxima occur at 161 and 598 Ma (Fig. 7). Tuffaceous sandstone sampled near the contact with overlying Eocene tuffs (ELKO-2; Fig. 2) yielded a maximum depositional age of 37.89 ± 0.46 Ma based upon inclusion of most measured grains in the sample (Fig. 6D). Only four older zircons were present in the sample (Fig. 7). Our new maximum depositional age constraints for the Elko Formation show that its deposition ranged from ca. 45.9 to 37.9 Ma (Table 2), in excellent agreement with prior work showing that the Elko Formation was deposited between ca. 46.1 and 38.9 Ma (Haynes, 2003). These results show that deposition of the Elko Formation spans a greater age range than assumed by Horton et al. (2004; 42–39.4 Ma), Mix et al. (2011; 42.6–39.4[?]Ma), and Mulch et al. (2015; 42.5–38.6[?] Ma), and the studies that cited these values (Chamberlain et al., 2012; Feng et al., 2013).
The lower part of the Eocene Elko Formation (sample ELKO-1; Fig. 2) records an abrupt shift in provenance (Fig. 7) expressed by the appearance of Eocene zircon. The oldest strata contain abundant ca. 46 Ma zircons (Figs. 6A and 7) whose age overlaps tightly with the age of volcanism within the Challis volcanic field (Sanford, 2005, and references therein), located north of the Snake River Plain in Idaho (Fig. 1A). The presence of Eocene volcanic material, including a tuff dated as ca. 46 Ma in the lower conglomerate member of the Elko Formation (Haynes, 2003; Table 2), differentiates the conglomerates of the Elko Formation from conglomerates in the older redbeds (TKcs) (Fig. 3B). Previous work throughout the northern Great Basin has documented reworking of Challis-derived volcanic detritus into local sedimentary basins (Haynes, 2003; Henry, 2008). Results from sample ELKO-3 exhibit a slightly younger ca. 45–44 Ma age peak (Figs. 6B and 7) at the young end of ages reported for Challis volcanism (Janecke and Snee, 1993; Janecke et al., 1999; Gaschnig et al., 2009; Chetel et al., 2011; Fig. 1A). Volcanism in the more proximal northeast Nevada volcanic field occurred between 42.6 and 39.0 Ma (Brooks et al., 1995; Fig. 1A). It is interesting that detrital material in this younger age range is absent in the Elko Formation, except in an area significantly to the northwest of the mapped area (Haynes, 2003). Detrital zircon peaks ca. 38 Ma are present at the very top of the Elko Formation within the mapped area (sample ELKO-2; Figs. 6C and 7), where they generally overwhelm other detrital zircon populations. These populations are interpreted as locally derived and to represent the onset of volcanism in the Robinson Mountain volcanic field ca. 38.5 Ma, continuing until 36.8 Ma (Table 2). Two zircons analyzed ca. 52 Ma in the lowermost Elko Formation (sample ELKO-1; Fig. 7) are slightly older than the Challis volcanic field of Idaho and southwest Montana, but are similar in age to the Lowland Creek volcanics and the early Dillon volcanic rocks of southwest Montana (Fritz et al., 2007; Gaschnig et al., 2009; Dudás et al., 2010; Chetel et al., 2011).
The age distributions of pre-Cenozoic zircons in the Elko Formation are broadly similar to those of the underlying Late Cretaceous(?)–Eocene(?) redbeds (Fig. 7). Lower Elko Formation deposits probably contain recycled zircons from the older redbeds and/or received recycled zircons from underlying Triassic and/or late Paleozoic passive margin sediments (Fig. 8). Unlike the older redbeds, Elko Formation samples ELKO-1, ELKO-3, and ELKO-2 all exhibit a Jurassic peak ca. 164–155 Ma (Fig. 7). Although Jurassic plutons are mapped in the RMEH and could represent a source for Jurassic detritus (e.g., du Bray, 2007), the Elko Formation lacks associated detrital muscovite or metamorphic lithics and other lithic fragments that would likely be derived from the uplift of deep-seated rocks in the RMEH. Based on these observations, it is unlikely that the RMEH was a source of sediment for the Elko Basin during the Eocene. It is most likely that the Jurassic zircon contained within the Elko Formation originated from the erosion of Jurassic volcanic and high-level intrusive rocks exposed in the nearby Cortez Mountains to the west of the Piñon Range and/or the more distant Delcer Buttes and Medicine Range east of the RMEH (e.g., Smith and Ketner, 1976, 1978; Ressel and Henry, 2006; Crafford, 2007; du Bray, 2007; Colgan and Henry, 2009). Smith and Ketner (1978) mapped a volcanic unit of possible Jurassic age (the Frenchie Creek Rhyolite) in the southeastern Piñon Range, southwest of the mapped area (Fig. 1B), but the eruption age of this unit is very uncertain (Smith and Ketner, 1976). The age of the oldest volcanic rocks erupted across the region of deposition of the Elko Formation provides a well-defined upper age bracket for the formation. Sample H10–45, collected from the lowest exposed cooling unit in the tuff of Robinson Mountain, yielded a 40Ar/39Ar age of 38.48 ± 0.15 Ma (C.D. Henry, 2011, personal commun.; Figs. 2 and 3B; Table 2).
Eocene Robinson Mountain Volcanic Field and Associated Sedimentary Rocks
Eocene volcanic, subvolcanic, and associated sedimentary rocks of the Robinson Mountain volcanic field were previously mapped as part of the Eocene–Oligocene Indian Well Formation (Smith and Ketner, 1976, 1978). However, new geochronology, together with geologic mapping, demonstrates that most sedimentary rocks in the map area originally assigned to the Indian Well Formation are instead part of the Miocene Humboldt Formation. We therefore suggest discontinuing the name Indian Well Formation, and simply refer to the Eocene igneous rocks on the basis of their presence within the Robinson Mountain volcanic field (Ressel and Henry, 2006). The stratigraphy of the Robinson Mountain volcanic field and its minor associated sedimentary rocks is described in the following.
Tuff of Cissilini Canyon and Associated Igneous and Sedimentary Rocks
The tuff of Cissilini Canyon (Ttcc) appears to have been deposited unconformably above the basal redbeds (TKcs) in the northwest part of the mapped area (Figs. 2 and 3). This tuff consists of smoky quartz–bearing felsic ignimbrites ranging from distinctive dark gray lithic-crystal tuffs with light gray pumice to biotite-rich light gray–pink crystal-vitric tuffs containing light gray–pink pumice. North of the map area, near Emigrant Spring (Fig. 1B), a possibly correlative white, biotite-rich rhyodacite tuff has yielded biotite and hornblende 40Ar/39Ar ages of 37.47 ± 0.11 and 38.0 ± 0.3 Ma, respectively (sample NEP-14 of R. Fleck, inHaynes, 2003; Table 2). Subvolcanic intrusions and flow-banded rhyolite flows and domes are exposed nearby (Fig. 2). Eocene intrusions in the vicinity of Emigrant Spring and the Railroad mining district (Fig. 1B) are dated between ca. 39.1 and 37.4 Ma (Ressel and Henry, 2006; Table 2).
Tuff of Robinson Mountain and Associated Igneous Rocks
The silicic tuff of Robinson Mountain (Ttrm) conformably overlies the Elko Formation (Fig. 2). It consists of heterogeneous, mostly moderately to densely welded lithic to crystal ignimbrites best exposed south of Robinson Mountain (Figs. 1B, 2, and 3C). This tuff is differentiated by tabular zoned plagioclase phenocrysts that are distinct from the clumped and concentrically zoned plagioclase observed in other Robinson Mountain volcanic rocks. A sanidine 40Ar/39Ar age of 38.48 ± 0.15 Ma (sample H10–45) was previously obtained from the lowermost flow, and another sanidine 40Ar/39Ar age of 37.70 ± 0.06 Ma (sample H10–47) was obtained from an upper flow (C.D. Henry, 2014, personal commun.). These ages both closely overlap with our maximum depositional age constraints for the underlying uppermost Elko Formation (Fig. 3B; Table 2).
Tuff of Dixie Creek and Associated Igneous and Sedimentary Rocks
The light gray crystal-vitric tuff of Dixie Creek (Ttdc) records at least 12 eruptive events that produced a poorly to moderately welded stack of calc-alkaline trachydacite ignimbrites (Figs. 2, 3D, and 3E). Although these and other nearby tuffs have been referred to as the “tuff of Jiggs” (Palmer et al., 1991; Gordee et al., 2000), their proximity to Dixie Creek, not the town of Jiggs (Fig. 1B), was the reason we informally renamed them. Weathering of the tuff below its more resistant upper part forms distinctive light gray rounded boulders.
Ion probe U-Pb zircon results indicate that the tuff of Dixie Creek was erupted between ca. 37.8 ± 0.4 and 37.3 ± 0.3 Ma (Table 2). Sample 10JLS06A consists of pumice collected from the third-highest exposed cooling event (Figs. 2 and 3). Of 12 zircons, 5 yield a maximum concordia intercept age of 37.7 ± 0.4 Ma (Fig. 4A); 4 additional grains constrain a slightly younger concordia intercept age of 36.5 ± 0.4 Ma, but this is less robust due to the relatively large uncertainties of the results used to constrain the intercept age (Fig. 4A). Sample 10JLS05 is pumice from the uppermost exposed cooling event (Figs. 2 and 3), and 5 of 12 zircons yield an intercept age of 37.34 ± 0.33 Ma (Table 2; Fig. 4B). These ages are consistent with a previously obtained sanidine 40Ar/39Ar age of 37.70 ± 0.38 Ma collected by D.A. John from one of the lowest flows exposed (sample 99-DJ-40; Ressel and Henry, 2006; Table 2).
Despite their similar appearance and age, there is evidence that the chemistry of the magma chamber these tuffs erupted from was heterogeneous or that it evolved during its eruptive history (Supplemental File). While the dominant white–light gray–pink pumice represented by samples 10JLS205 and 10JLS206A is dacite in composition, less common dacite with a papery texture (12JLS211) and more mafic trachydacite to trachyandesite pumice (12JLS210) are also present (Figs. A3 and A4 in the Supplemental File). The strong resorption textures exhibited by pumice with different geochemistries in single eruption units (Supplemental File) suggests mixing of more mafic magma with the dominant dacitic composition magma at the time of eruption.
Based on the geochemistry of vitrophyres, nearby lava flows and domes consist of both peraluminous and calc-alkaline rhyolite (unit Tr, Figs. 2 and 3; Figs. A3 and A4 in the Supplemental File) that were probably erupted in multiple stages because they both locally underlie and overlie the tuff of Dixie Creek. Rhyolites range from glassy and crystal poor to phenocryst rich, and are transitional in some places to subvolcanic intrusions (see following discussion). Most flows are mantled above and/or below by vitrophyres and breccias, allowing them to be separated into individual cooling units. The rhyolite flows and domes share numerous petrological characteristics with flow-banded 35.1 ± 0.5 Ma dacite lava domes exposed in the Sulphur Spring Range (unit Tpd of Ryskamp et al., 2008; Fig. 1B), including similar large plagioclase phenocrysts.
Subvolcanic Intrusions and Associated Sedimentary Deposits
Weakly calc-alkaline rhyolite porphyries form the tall edifices of Robinson Mountain and Squaw Mountain, which were probably large volcanic necks that intruded local ignimbrites (Figs. 1B and 2; Supplemental File). This most common type of subvolcanic intrusion (Tsvi) typically contains complex clusters of large plagioclase phenocrysts that are commonly concentrically zoned and closely resemble those in the rhyolite flows and domes and the tuff of Dixie Creek. These and other petrologic similarities suggest a shared magmatic source that differs somewhat from the otherwise compositionally similar tuff of Robinson Mountain, which shows simply zoned, monocrystalline plagioclase phenocrysts. Subvolcanic intrusions having different mineralogy or appearance from the dominant type are exposed locally in parts of the map area (Tsviw and Tsvip, Fig. 2).
Subvolcanic intrusions mostly postdate and intrude the Eocene volcanic rocks, including the ignimbrites and rhyolite domes (Fig. 2). However, in one location west of the southern part of Cedar Ridge, a tabular subvolcanic intrusion may represent a partially exhumed conduit that fed eruption of nearby rhyolite flows and domes (Fig. 2). A sanidine 40Ar/39Ar age of 37.51 ± 0.38 Ma was previously obtained from this intrusion (sample 99-DJ-31; Ressel and Henry, 2006; Table 2). This result is similar to our new ion microprobe U-Pb zircon crystallization age of 37.64 ± 0.74 Ma (sample 11JLS104) for an exposure of the Tsvi unit elsewhere in the mapped area (Fig. 2; Table 2). The crystallization age for sample 11JLS104 is based on results obtained from 4 of 16 grains (Figs. 2 and 4C). We interpret older grains detected in 11JLS104 as inherited from the source region. We obtained a younger ion microprobe U-Pb zircon crystallization age of 36.84 ± 0.34 Ma for an additional sample (10JLS11A). We used 4 of 14 zircon results to constrain the age for sample 11JLS11A, excluding a single, discordant measurement with a high error (Fig. 4E). A concordia upper intercept age of 2467.8 ± 6.3 Ma was obtained for this sample (Fig. 4D).
Thin sandstone, siltstone, and shale (unit Ts) layers are commonly exposed adjacent to subvolcanic intrusions or rhyolite flows and domes (Fig. 2), which could represent detrital sources for the sedimentary layers. These sedimentary deposits typically contain clasts of felsic volcanic rocks resembling Robinson Mountain volcanic rocks. A U-Pb detrital zircon sample, 12JLS161, was collected from a possibly reworked or mass-wasted volcanic rock in one such exposure; this established a maximum depositional age of 37.67 ± 0.27 Ma (Fig. 7; Table 2), based on 22 of 83 measurements (Fig. 6D). One younger zircon analysis (ca. 33.3 ± 1.1 Ma; Table A5 in the Supplemental File) was not considered because a single analysis is not sufficient to establish a maximum depositional age, but it is possible that the depositional age of sample 12JLS161 is younger than ca. 33 Ma (Fig. 6D). The large population of ca. 37.7 Ma detrital zircon grains (Fig. 6D) in sample 12JLS161 was likely derived from the contemporaneous (Table 2) tuff of Dixie Creek or the nearby subvolcanic intrusions (Fig. 2). Although 12 analyses yielded Proterozoic ages with peaks at 1.73 Ga, 1.41 Ga, and 1.26 Ga, cathodoluminescence imaging of these grains reveals that most analyses were obtained from the cores of igneous zircons (Figs. 7 and 9). The core ages match age maxima prevalent in Devonian, Cambrian, and Neoproterozoic strata of the Great Basin region (Gehrels and Pecha, 2014). We thus interpret the zircon cores to represent material inherited from either Precambrian basement and/or Neoproterozoic to early Paleozoic strata in the source region of late Eocene magma genesis deep in the crust.
A second sedimentary sample in contact with Eocene volcanic and subvolcanic rocks (TIWB-1) is characterized by bedded, biotite-rich, white sandstones, and yielded an Eocene maximum depositional age (Fig. 6E) of 37.33 ± 0.14 Ma (on the basis of 67 of 94 analyses), nearly identical to ages obtained for nearby subvolcanic intrusions (Table 2). There is also a small age maximum ca. 159 Ma (Fig. 7). Applying the same conservative approach that was used with 12JLS161, we neglected 3 slightly younger U-Pb ages in specimen TIWB-1 that could be geologically meaningful. Based upon the youngest of these ages, the depositional age of TIWB-1 could be younger than ca. 35 Ma (Fig. 6E). A Jurassic peak at 159 Ma indicates either continued detrital zircon input from nearby Jurassic high-level volcanic and intrusive rocks or recycling of older Cenozoic strata containing material of Jurassic age (Fig. 7).
Oligocene(?) Volcanic and Associated Sedimentary Rocks
Mafic to Intermediate Composition Volcanic Rocks
At least two flows of platy, vesicular, dark gray, and nearly aphanitic basaltic andesite (Tba) with a total thickness of >150 m overlie rhyolite flows of the Eocene Robinson Mountain volcanic field (Figs. 2 and 3D). In addition, a small (≤0.2 km2 map area), black, isolated basaltic trachyandesite intrusion (sample 12JLS125, Supplemental File) is exposed to the west of an east-dipping normal fault system that bounds the east side of the Piñon Range (Figs. 2 and 3D, unit Tbta). For discussions of additional unit descriptions and speculative age constraints, see Lund Snee (2013) and Lund Snee and Miller (2015).
Sedimentary Rocks and the Oligocene Tuff of Hackwood Ranch
Shallowly west dipping, thin (∼35 m), poorly consolidated, cross-bedded, silty sandstone and pebble conglomerate were deposited above the tuff of Dixie Creek, across a 10°–15° angular unconformity that dips to the west (Palmer et al., 1991; Gordee et al., 2000). Mapped in Figure 2 as the stratigraphically lowest part of unit Tthr, these rocks have the same dip angle as the overlying Oligocene tuff of Hackwood Ranch (unit Tthr) and demonstrate that underlying Eocene volcanic rocks were tilted before deposition of the Oligocene sediments. Additional westward tilting occurred after eruption and deposition of the tuff of Hackwood Ranch.
Detrital zircon results from a sample near the stratigraphic middle of this sedimentary deposit indicate a maximum depositional age of 33.86 ± 0.40 Ma (sample 10JLS08, Fig. 6F; Table 2), based on 18 of 28 grains. Most of the older zircon in this sample (ca. 41–36 Ma) is similar in age to the Robinson Mountain volcanic rocks, and a single Cretaceous zircon is also present. No igneous rocks with ages close to the main 33.9 Ma zircon population in sample 10JLS08 are present in the map area (Table 2). Possible sources for these zircons include the ca. 33.8 Ma Caetano Tuff or the older and less widespread 34.2 Ma tuff of Cove Mine, which also erupted from the Caetano caldera to the southwest (John et al., 2008; Fig. 1A).
The tuff of Hackwood Ranch consists of three cliff-forming, poorly welded eruptive units (Figs. 2 and 3E). On the basis of phenocryst content, the tuff of Hackwood Ranch is transitional between a quartz latite and latite; however, resorption textures in some phenocrysts suggest that they are xenocrystic. Whole-rock (11JLS105) and pumice-only (HACK-1) geochemical analyses indicate that the tuff is a high-silica rhyolite (Supplemental File). Outcrop and thin section evidence for posteruptive silicification suggests that its high silica (and low alkali) content is due to secondary alteration. The age of the tuff of Hackwood Ranch places constraints on the minimum depositional age of the underlying sedimentary deposits. Ion probe U-Pb ages of zircon and 40Ar/39Ar sanidine ages for samples of the tuff of Hackwood Ranch range from 31.59 to 31.08 Ma (Table 2), with 31.08 Ma representing our interpretation of the best age for its eruption. Zircon extracted from pumice in sample HACK-1 yielded a maximum eruptive age of 31.59 ± 0.29 Ma (12 of 16 grains; Fig. 4F). Single crystal, total fusion 40Ar/39Ar ages of sanidine from sample HACK-1 yielded a weighted mean age of 31.10 ± 0.47 Ma (2σ total error; mean square of weighted deviates, MSWD = 0.35). Of 14 grains, 2 were rejected as detrital contamination. Zircon recovered from pumice and ash matrix from the second sample (11JLS105) yielded a maximum eruptive age of 31.26 ± 0.30 Ma (8 of 16 grains; Fig. 4G). The second sample (11JLS105) yielded a 40Ar/39Ar weighted mean age of 31.08 ± 0.47 Ma (10 of 13 grains; Figs. 3E and 5A; Table 2). The agreement of the 40Ar/39Ar and U-Pb age results within 95% confidence from 2 samples of the tuff of Hackwood Ranch strengthens our conclusion that the eruptive age of the tuff of Hackwood Ranch was 31.08 Ma.
Although the age of the Hackwood Ranch tuff overlaps with zircon and monazite ages from metamorphic and igneous rocks in the northern RMEH (as young as 29 Ma; Wright and Snoke, 1993; MacCready et al., 1997; Howard et al., 2011; Fig. 10), this tuff is not thought to have been erupted locally, but most likely erupted from a caldera near the Desatoya and Clan Alpine Mountains in central Nevada (Fig. 1A). Some tuffs traveled hundreds of kilometers from their eruptive centers, and the 31.2 Ma tuff of Rattlesnake Canyon is closest in age to the tuff of Hackwood Ranch (Faulds et al., 2005; Henry et al., 2012; Henry and John, 2013). A similar age of 31.40 +1.3/–0.5 Ma was also obtained from U-Pb LA-ICP-MS analysis of zircon from an andesite in the Sulphur Spring Range to the southwest (sample 04EB123, Ryskamp et al., 2008; Table 2). Otherwise, this is an unusual age for volcanic rocks in northern Nevada, which erupted across a very narrow time interval (Fig. 10).
Miocene Humboldt Formation
The name Humboldt Formation (Th) has been applied to diverse sedimentary rocks of Miocene age deposited in basins in northeastern Nevada (Sharp, 1939; Smith and Ketner, 1976, 1978; Smith and Howard, 1977; Solomon et al., 1979; Stewart, 1980; Wallace et al., 2008; Fig. 1B). Sharp (1939) initially measured ∼1800 m of Humboldt Formation along the eastern flanks of Huntington Creek, west of the RMEH (Fig. 1B). Smith and Ketner (1978) restricted the name Humboldt Formation to the middle and part of the upper members of the Sharp (1939) section, measuring >560 m thick and consisting of upper Miocene conglomerate, sandstone, siltstone, and claystone with beds of limestone, tuff, and ash. West of these deposits, Smith and Ketner (1978) mapped a large portion of the present study area as latest Oligocene–Eocene (time scale of Walker et al., 2012) Indian Well Formation (Fig. 1B), and subsequent workers also used this name (Solomon et al., 1979; Satarugsa and Johnson, 2000; Haynes, 2003; Horton et al., 2004; Chamberlain et al., 2012; Mulch et al., 2015).
Geologic mapping with supplemental geochronology (this study) has determined that most of the sediments previously mapped as Indian Well Formation west of Cedar Ridge and Red Spring and near Indian Well (including the type section of Smith and Ketner, 1976), are significantly younger strata belonging to the mostly Miocene (and possibly as old as latest Oligocene) Humboldt Formation (Figs. 1B and 2). If we include these strata in the Humboldt Formation, its total thickness is ∼2130 m in the mapped area alone (Figs. 2 and 3E).
Tephra is present within Miocene sections in northeast Nevada beginning ca. 16 Ma, derived from the explosive rhyolitic volcanic centers of the Snake River Plain–Yellowstone hotspot track (Perkins et al., 1998; Perkins and Nash, 2002; Wallace et al., 2008; Colgan et al., 2010; Fig. 1A). The rhyolites were anhydrous and erupted at high temperature, and they contain little or no biotite (Honjo et al., 1992; Perkins and Nash, 2002; Cathey and Nash, 2009; Ellis et al., 2010; Konstantinou et al., 2012). Snake River Plain tephra thus provide a good means for establishing maximum depositional ages and the stratigraphic age range of the Humboldt Formation. Samples analyzed from the Humboldt Formation reveal at least 5 Miocene age populations: ca. 15.6 Ma, ca. 14.2 Ma, ca. 12.4 Ma, ca. 10.9 Ma, and ca. 8.2 Ma (Figs. 5–7). The 15.6 Ma peak could be derived from several Snake River Plain volcanic sources in Oregon and Nevada, including the High Rock and McDermitt caldera complexes or the Lake Owyhee volcanic field, and the ca. 14.2 Ma tephra could be sourced from the Owyhee-Humboldt volcanic field on the Idaho-Utah-Nevada border (Perkins et al., 1998; Perkins and Nash, 2002; Coble and Mahood, 2012; Fig. 1A). Zircons that are 12.4 Ma in age could be derived from eruptions at the Bruneau-Jarbridge volcanic field of southern Idaho, and ca. 10.9 Ma and ca. 8.2 Ma zircon ages may represent air fall from the Bruneau-Jarbridge, Twin Falls, or Picabo volcanic fields (Perkins and Nash, 2002; Cathey and Nash, 2009; Ellis et al., 2010; Konstantinou et al., 2012). Rare zircons of ca. 32–31 Ma age are also present in Miocene sedimentary rocks and might be derived from the erosion of the underlying tuff of Hackwood Ranch.
Based on map relationships (Fig. 2) and the above ages, the exposures of the Humboldt Formation are youngest eastward toward Huntington Valley (Colgan et al., 2010; this study). The oldest strata of the Humboldt Formation yield detrital zircon age distributions dominated by latest Oligocene to middle Miocene zircon with essentially no zircon older than 200 Ma (see samples ELM11-PN19 and Tiws-J4 in Fig. 7). The Humboldt Formation has been repeated by gentle folds and a west-dipping normal fault system with at least 700 m of throw near Cedar Ridge and Red Spring (herein called the Cedar Ridge fault system) (Wallace et al., 2008; Figs. 1B, 2, and 3E). West of the Cedar Ridge fault system, strata of the Humboldt Formation become finer and more carbonate rich upsection, suggesting a transition from fluvial to lacustrine depositional environments. However, east of the fault system the percentage and thickness of pebble conglomerate horizons increases stratigraphically upward, suggesting a depositional environment increasingly dominated by fluvial activity, consistent with the interpretation of Sharp (1939). Dense, dark blue-gray chert exposed near the normal fault system (unit Ths, Fig. 2) is interpreted to be silicified Humboldt Formation related to hydrothermal activity surrounding the fault system or to silicification in a lacustrine setting. East of the fault system, the upper parts of the mostly Miocene Humboldt Formation (samples 12HBD05 and 12HBD09) appear to represent an unroofing sequence in which previously deeply buried strata of the RMEH were eroded as they were exhumed by middle Miocene faulting (e.g., Colgan et al., 2010).
The basal beds of the Humboldt Formation overlie the Oligocene tuff of Hackwood Ranch across a ≥30° angular unconformity that is mapped in float but is not directly exposed (Figs. 2 and 3E). These lowest deposits include poorly sorted, cross-bedded pebble conglomerate, sandstone, and siltstone with interbedded volcanic ash and clasts resembling Eocene igneous rocks and Paleozoic metasedimentary rocks exposed nearby (Figs. 2 and 3E). Sample ELM11-PN19 is sandstone that was sampled from close to the stratigraphically lowest exposures of the Humboldt Formation (Figs. 2 and 3E). It is volcanogenic and yielded a U-Pb detrital zircon maximum depositional age of 24.39 ± 0.08 Ma on the basis of 90 of 98 grains (Fig. 6G). In calculating this result, we excluded 2 slightly younger ages of 23.3 ± 0.8 Ma and 23.5 ± 0.7 Ma. Possible sources for this mostly ca. 24.4 Ma detrital zircon population include the Fish Creek Mountains Tuff from the Fish Creek Mountains caldera west of the Shoshone Range, that erupted ca. 24.7 Ma (John et al., 2008), or 25–24 Ma eruptions at the Elevenmile Canyon and Poco Canyon calderas of the Stillwater Range to the west (John, 1995; Fig. 1A).
Detrital zircon sample TIWS-J4 was collected from ∼250 m above the exposed base of the section (Fig. 3E) and represents the sandstone matrix of volcanic ash–rich, cross-bedded, pebble conglomerate within the Humboldt Formation. This sample yielded a significantly younger U-Pb detrital zircon maximum depositional age of 15.64 ± 0.11 Ma (Fig. 6H; Table 2), based on 42 of 91 analyses. No younger grains were excluded. This upper bound for the depositional age is in good agreement with the 40Ar/39Ar sanidine constraints (Table 2) obtained from the lowest air fall tuff exposed in the Humboldt Formation (see following). Sample TIWS-J4 also contains older Cenozoic and a few Cretaceous zircons (Fig. 7). Our sample TIWS-J4 was collected near the locality where several samples collected by Chamberlain et al. (2012) and Mulch et al. (2015) have depositional ages spanning 38.9–38.0 Ma based on 40Ar/39Ar dating of biotite and assumed sedimentation rates (NV-IW 01–07 through NV-IW 20–07; ages and sample numbers slightly different in Mulch et al., 2015; Figs. 2 and 3E). Because our geochronologic analyses yielded many zircon grains with significantly younger (Miocene) ages in the same part of the stratigraphic section dated by them, we interpret their ages as dating detrital biotites eroded from the underlying Eocene Robinson Mountain volcanic field, which is exposed nearby, is extremely rich in biotite, and is of the same (ca. 38 Ma) age range (discussed in preceding). In contrast, high-temperature, anhydrous eruptions from the Snake River Plain produced very biotite-poor tephra that was deposited in some of the same strata during Miocene time (Perkins et al., 1998; Wallace et al., 2008; Colgan et al., 2010; Ellis et al., 2010). As the Humboldt basin formed, the Robinson Mountain volcanic rocks were tilted to the east and debris from their erosion was delivered into the basin from its western side (Fig. 2). Boulder to cobble horizons derived almost exclusively from these volcanic rocks are common in the western exposures of the Humboldt Formation.
Above its lowest exposures of conglomeratic sandstone, the Humboldt Formation is characterized by marl with local limestone and calcite-cemented conglomerate beds and horizons of poorly consolidated, ∼5–20-m-thick, light tan to blue-gray Snake River Plain volcanic ash (Smith and Ketner, 1976; this study; Figs. 2 and 3E). Alkali feldspar 40Ar/39Ar ages obtained from samples of air fall tuffs within the lower parts of the Humboldt Formation place additional bounds on the maximum depositional ages (Fig. 5). However, because these air fall tuffs may have been in part or entirely reworked, the ages obtained from them could be somewhat older than their true depositional age. Sample ELM11-PN13, collected from the lowest Snake River Plain air fall tuff exposed in the Humboldt Formation (Figs. 2 and 3E), yielded an alkali feldspar 40Ar/39Ar maximum eruptive age of 15.78 ± 0.25 Ma (10 of 11 grains; Fig. 5C). Sample ELM11-PN11 (Figs. 2 and 3E), collected from an along-strike exposure of the same air fall tuff as ELM11-PN13, yielded an alkali feldspar 40Ar/39Ar maximum eruptive age of 15.51 ± 0.24 Ma (8 of 11 grains; Fig. 5D). Although 40Ar/39Ar ages of ELM11-PN13 and ELM11-PN11 overlap at ± 2σ total error, it is problematic that their ages fail to overlap at ± 2σ analytical error because these samples were coirradiated and collected from the same air fall tuff (15.78 ± 0.07 Ma; MSWD = 1.42 for ELM11-PN13 versus 15.51 ± 0.03; MSWD = 1.40 for ELM11-PN11; Table A1 in the Supplemental File). Variability in Ca/K ratios and age expressed by high MSWD values (Figs. A1C and A1D in the Supplemental File) suggests that our results may have been affected by detrital contamination that occurred during reworking of the ash deposits by surficial processes. Sample ELM11-PN11 was collected at roughly the same locality as Indian Well Formation samples NV-IW 32–07 through NV-IW 40–07 of Chamberlain et al. (2012) and Mulch et al. (2015); their reported ages span the Eocene to earliest Miocene (38.0–22.3 Ma) and are based on 40Ar/39Ar and K-Ar of biotite and sanidine, as well as assumed sedimentation rates between dated samples. As discussed herein, our new data clearly indicate a middle Miocene age for this part of the Humboldt Formation, based on multiple U-Pb and 40Ar/39Ar data sets. We interpret the results of Chamberlain et al. (2012) and Mulch et al. (2015) as ages obtained on detrital biotite derived from the underlying Robinson Mountain volcanic field, as this biotite-rich debris is extremely common in this part of the Humboldt Formation (Fig. 3E).
Sample TIWS-J3 was collected from a stratigraphically higher air fall tuff in the Humboldt Formation (Fig. 2). We calculated a weighted mean age of 14.62 ± 0.22 Ma from the youngest 7 of 20 grains in this sample (Figs. 3E and 5E). However, the high MSWD (7.1) associated with this result indicates that our selection is unlikely to define a homogeneous age distribution. We thus regard the TIWS-J3 result as a maximum bound upon the depositional age of the strata that overlie the tuff.
Horton et al. (2004) and Chamberlain et al. (2012) assigned pre-Miocene ages to the section of the Miocene Humboldt Formation that is below the locality of sample TIWS-J3 (Figs. 1B, 2, and 3E), which was previously mapped as the Indian Well Formation. Using the assumption of constant sediment accumulation rates between (and extrapolated beyond) their samples dated by the 40Ar/39Ar method on detrital biotite, the Indian Well Formation was believed by Horton et al. (2004) and Chamberlain et al. (2012) to span ca. 38.9–22.3 Ma; our reassignment of these strata to the mostly Miocene Humboldt Formation is significant because they reported a major isotopic shift (∼6‰ δ18O decrease) near the contact between the Elko Formation and the overlying Indian Well Formation, which they reported was constrained between ca. 40 and 38 Ma (Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Feng et al., 2013; Mulch et al., 2015). This shift is cited as evidence for a rapid elevation gain (∼2.5 km in <2 m.y.) (Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Feng et al., 2013) related to volcanism reaching this latitude at 38 Ma. Previous workers (Chamberlain et al., 2012) inferred middle Eocene topographic relief of as much as 2.2 km in northeast Nevada by subtracting lower and upper bound elevation estimates determined within this general region. The lower bound elevation estimates of 2.0 ± 0.2 km in the late-middle Eocene (ca. 42–39 Ma) were determined using leaf physiognomic methods for strata collected in Copper Basin, Nevada (Wolfe et al., 1998). The upper bound elevation estimates of as much as ∼4.2 km hypsometric mean (Mix et al., 2011; Chamberlain et al., 2012) were obtained by stable isotope methods for strata from Huntington Valley, ∼130 km south of Copper Basin (Fig. 1A), since the strata in question were originally thought to be of similar age. Because all samples collected and dated from Huntington Valley that were originally thought to span ca. 38.9–22.3 Ma are now known to be ca. 16–15 Ma, the upper bound paleoelevation postulated for northeast Nevada across this time span instead probably represents conditions in the middle Miocene (ca. 16–15 Ma), not the Eocene.
Sample ELM11-PN2 was collected from the stratigraphically highest air fall tuff exposed in the section of the Humboldt Formation exposed west of the Cedar Ridge fault system (Figs. 2 and 3E). The sample yielded sanidines having a weighted mean age of 12.35 ± 0.19 Ma and a MSWD of 1.90 for 13 of 13 grains (Fig. 5F). The Humboldt Formation is partially repeated east of the Cedar Ridge normal fault system (Wallace et al., 2008; Figs. 2 and 3E). The lowest mapped stratigraphic levels on the east side of the fault system unconformably overlie an undated unit resembling the Eocene Elko Formation (Figs. 1B and 2). A single detrital zircon grain from pebble conglomerate interpreted to represent low stratigraphic levels of the Humboldt Formation east of the Cedar Ridge fault system (sample 12HBD06; Fig. 2) yielded a poorly constrained maximum depositional age of 15.7 ± 0.5 Ma. However, unlike samples ELM11-PN19 and TIWS-J4, this sample is volcanic-fragment poor, yielding only one post-Eocene zircon analysis (15.7 ± 0.5 Ma) and three significantly older analyses forming a small maximum ca. 38 Ma (Fig. 7). The majority of zircon in this sample defines age peaks at 1.09 Ga, 1.30 Ga, 1.52 Ga, and 1.60 Ga, with a smaller maximum at 421 Ma (Fig. 7). It is interesting that a large proportion (>40%) of grains analyzed in 12HBD06 have high 204Pb (>400 counts per second; many are >1000 counts per second; Table A5 in the Supplemental File). This might be related to hydrothermal fluid circulation around the Cedar Ridge fault system, which could have introduced common Pb into older, radiation damaged zircon.
Sample 12HBD06 marks the transition from sediments with volcanic-dominated zircon to recycled zircon from sedimentary sources, as also evidenced by the high proportion of Paleozoic clasts in this sample. The age distribution of pre–200 Ma zircon extracted from 12HBD06 is similar to that from the Late Cretaceous(?)–Eocene(?) redbeds (TKcs) and the Eocene Elko Formation. This includes one 421 Ma analysis and populations at 1.15–0.90 Ga and 1.60–1.25 Ga. The similar age distributions (Fig. 7) of sample 12HBD06 and the Late Cretaceous–Eocene redbeds (TKcs) suggest that sediment contained within both was derived from erosion of Mesozoic and late Paleozoic strata exposed throughout the northern Great Basin or recycled from sediments eroded from these formations.
Uppermost strata of the Humboldt Formation in the mapped area (Figs. 2 and 3E) are light tan, poorly consolidated marls, siltstones, sandstones, and pebbly sandstones with interbedded vitric tuffs. In contrast to 12HBD06, samples collected from the youngest portion of the Humboldt Formation (12HBD05 and 12HBD09) contain clastic material that can be more confidently inferred to have been derived from the erosion of metamorphic rocks in the RMEH. Detrital muscovite, epidote, and lithic fragments of strained quartzite are observed in some of the stratigraphically high samples of the Humboldt Formation (12HBD05; Fig. 2), suggesting input from the erosion of the RMEH. Sample 12HBD05 is Humboldt Formation sandstone that yielded sufficient young zircon (9 of 92 grains) to define a maximum depositional age of 14.15 ± 0.21 Ma (Fig. 6I). This result is similar to 40Ar/39Ar results from volcanic tuff sample TIWS-J3 (Table 2), west of the Cedar Ridge fault system, suggesting that it may be stratigraphically equivalent. Well-defined U-Pb age maxima span the Paleoproterozoic through Paleozoic, and minor populations are detected at 159 Ma and between 91 and 55 Ma, with minor maxima at 71 and 61 Ma, and between 37 and 23 Ma (Fig. 7).
The presence of Jurassic and minor early Cenozoic to Late Cretaceous detrital zircon in this sandstone is consistent with derivation from eroded granites and pegmatites in the RMEH (Fig. 10; Howard et al., 2011). We consider it significant that Late Cretaceous detrital zircon was, in general, not detected in older samples of the Humboldt Formation that we have studied. Rare Late Cretaceous ages of ca. 74 Ma, ca. 90 Ma, and ca. 79 Ma were found in samples of Eocene strata (TIWS-J4, TIWB-1, and ELKO-1, respectively; Fig. 7), but those are solitary ages and are not accompanied by the presence of metamorphic clasts and muscovite. Therefore, we consider it highly unlikely that the rare Late Cretaceous ages in the underlying Eocene strata represent an influx of detrital material from the RMEH. Instead, the signal that we can confidently tie to the RMEH is detected only in the uppermost Humboldt Formation strata exposed in the mapped area. This finding is consistent with 87Sr/86Sr values that increase from lower values of 0.7076–0.7085 in Miocene Humboldt Formation strata within the study area that are now known to be ca. 15 Ma or older (Fig. 3E), and 0.7092 at the base of a section measured <5 km east of the mapped area (Fig. 1B), to 0.7106 stratigraphically higher within the Humboldt Formation (Mulch et al., 2015). This trend of increasing 87Sr/86Sr values upward in the Miocene Humboldt Formation supports our interpretation that previously deeply buried rocks of the RMEH were first exposed midway through Humboldt Formation deposition (ca. 14.2 Ma), thereafter providing a new basement-derived sediment source to the evolving basin.
The youngest maximum depositional age for the Humboldt Formation was obtained from sandstone sample 12HBD09 (Figs. 2 and 3E) using U-Pb detrital zircon geochronology. An age of 8.16 ± 0.15 Ma was obtained from 6 of 94 grains (Fig. 6J; Table 2). This result is the youngest depositional age constraint reported for the Humboldt Formation in Huntington Valley by >1 m.y. (Perkins et al., 1998; Wallace et al., 2008; Colgan et al., 2010). Like samples 12HBD05 and 12HBD06, sample 12HBD09 is poor in Cenozoic zircon. Age populations for this sample range from Miocene to Neoarchean, with major maxima at 374 Ma, 1.02 Ga, 1.45 Ga, 1.78 Ga, and 2.54 Ga (Fig. 7). Additional maxima in sample 12HBD09 are at 16 Ma, ca. 40–38 Ma, 42–41 Ma, and 160 Ma with sparse intervening ages, including 2 grains analyzed near 24 Ma and a single grain analyzed ca. 97 Ma (Fig. 7).
The presence of metamorphic clasts that include fragments of quartz mylonite, and the detrital zircon age maxima in sample 12HBD09, supports derivation of sediment from the RMEH. While the 42–38 Ma populations in this sample closely match similar populations in the Eocene Elko Formation, zircon of this age would also be explained by erosion of the RMEH because these ages partly overlap with plutonic ages in the RMEH, such as the 38 Ma Harrison Pass pluton (Colgan et al., 2010). Older quartz diorite and quartz dioritic orthogneisses in the RMEH range from ca. 40 to 29 Ma (Wright and Snoke, 1993). Another important feature of the detrital zircon age distribution of sample 12HBD09 from the uppermost Humboldt Formation is the nature of its Proterozoic age distribution. Detrital zircon populations at 1.12 Ga, 1.45 Ga, 1.70–2.04 Ga, and 2.69 Ga in this youngest Humboldt Formation sample (Fig. 7) match similar peaks in the Neoproterozoic–early Cambrian Geertsen Canyon Quartzite, early Cambrian Osgood Mountain Quartzite, and Middle Ordovician Eureka Quartzite (Gehrels and Pecha, 2014; Fig. 8), and indicate that erosion of early Paleozoic and late Neoproterozoic strata in the RMEH may have contributed debris to the upper Humboldt Formation. Note that the abundant 1.4 Ga and 1.7 Ga zircon prevalent in 12HBD09 (Fig. 7) does not closely match the Grenville-dominated Proterozoic detrital zircon age distributions that would be contributed by Permian and Triassic strata in the upper part of the passive margin succession (Fig. 8). We thus interpret the Mesoproterozoic zircon in sample 12HBD09 as recycled from the erosion of the structurally deepest, Neoproterozoic and early Paleozoic part of the miogeoclinal succession exposed in the RMEH. The results from sample 12HBD09 are consistent with the interpretation that Neoproterozoic and early Paleozoic rocks in the RMEH were not exposed until extensional faulting began in the middle Miocene (Colgan et al., 2010).
In summary, the Humboldt Formation documents an unroofing sequence produced by the progressive fault-related exhumation and erosion of the RMEH. Humboldt Formation strata recycled zircons eroded from late Paleozoic and Mesozoic sedimentary rocks exposed at the paleosurface beneath the Tertiary unconformity until ca. 14.2 Ma or earlier. A detrital zircon signature that suggests derivation from RMEH plutonic rocks is first detected in sample 12HBD05 (Figs. 2, 3E, and 7), collected from beds that also contain abundant detrital muscovite and metamorphic clasts, which yielded a maximum depositional age of 14.15 ± 0.21 Ma (Table 2). This was followed by the arrival, at 8.2 Ma or earlier (sample 12HBD09; Table 2), of Precambrian and Cambrian sediment sources, which are the lowest stratigraphic levels currently exposed in the southern RMEH (Willden and Kistler, 1979; Crafford, 2007). This documented sequence of events suggests that the faulting that brought these rocks to the surface occurred in the Miocene, consistent with recent work showing that rapid fault slip unroofed the southern RMEH between ca. 17–16 Ma and ca. 12–10 Ma (Colgan et al., 2010).
Post-Miocene Sedimentary Rocks
The Hay Ranch Formation (Regnier, 1960; Smith and Ketner, 1976) consists of poorly consolidated conglomerate, sandstone, and siltstone. Fossil and zircon fission-track dating indicate a maximum depositional age of middle Pliocene–middle Pleistocene (Smith and Ketner, 1976). In addition, Quaternary gravel deposits commonly cap hilltops and slopes, indicating that broad alluvial surfaces were once developed across the map area and have since been eroded, cut by streams, and filled with younger Quaternary deposits (Fig. 2). For more detailed descriptions of these units, see Lund Snee (2013) and Lund Snee and Miller (2015).
Upper Crustal Perspective on the Evolution of the RMEH MCC
This study utilizes geologic mapping, geochronology, and geochemistry to understand the Late Cretaceous through Cenozoic upper crustal geologic history adjacent to the southern and central parts of the developing RMEH MCC in order to address some of the controversies about the evolution of the core complex and its age and mechanism of exhumation. Our geochronologic data from supracrustal sections enable high-temporal-resolution comparison between events documented at the surface of the Earth to those occurring in the deeper crust, as represented by igneous and metamorphic rocks of the RMEH MCC (Fig. 10).
Throughout the well-documented history of Late Cretaceous to Oligocene metamorphism, partial melting, and lower crustal flow within the developing RMEH MCC (e.g., Wright and Snoke, 1993; McGrew and Snee, 1994; MacCready et al., 1997; McGrew et al., 2000; Howard et al., 2011), events at the Earth’s surface represent relative quiescence (Figs. 10 and 11). The oldest deposits in the mapped area are above a regional late Mesozoic or basal Cenozoic unconformity and include redbeds, conglomerate, sandstone, siltstone, and limestone of Late Cretaceous(?)–Eocene(?) age representing nonmarine continental and lacustrine depositional settings. Their detrital zircon signature indicates recycling from local underlying bedrock units and/or their regional correlatives. It remains unclear whether deposition was continuous between these basal Cenozoic strata and the overlying Eocene Elko Formation, because the contact between the two units is not well exposed in the study area, but others have suggested that the Elko Formation may unconformably overlie the underlying TKl limestones and that there may be slight angular discordance within TKl and TKcs (Smith and Ketner, 1976). Deposition of the Elko Formation began ca. 46 Ma (Figs. 10 and 11), as recorded by the presence of a 45.9 ± 1.0 Ma tuff near its base (Haynes, 2003; this study; Table 2). This tuff, and another dated as 45.0 ± 0.5 Ma (Table 2), were likely erupted from the Challis volcanic field, the closest region with eruptive units of that age (e.g., Gaschnig et al., 2009; Fig. 1A). The onset of Eocene volcanism in Idaho and Montana was coeval with MCC development north of the Snake River Plain and involved the development of topographically elevated regions that fed major river systems that delivered sediment to the Pacific coast (Dumitru et al., 2013). Thus it is possible that rivers could also have drained southward from Idaho and Montana at this time, transporting material into the Elko Basin. This would suggest higher topography to the north than to the south during the Eocene. Gans (1990), Mix et al. (2011), and Chamberlain et al. (2012) suggested that a southward-propagating, thermally driven topographic high or bulge could have accompanied the well-documented southward-sweeping onset of volcanism. If so, the Elko Basin could have developed as a topographic low relative to areas further north that were elevated by the onset of volcanism at 42–38 Ma (Brooks et al., 1995; Ressel and Henry, 2006; Konstantinou et al., 2012; this study). However, the present-day extent of the Elko Basin is elongated in the north-south direction (Haynes, 2003), and this is particularly true after removal of Cenozoic extension, so relating its development to topographic effects associated with southward-sweeping volcanism is speculative. Although the Elko Formation is more spatially continuous than the underlying TKcs and TKl units, it is <180 m thick in the mapped area (Figs. 2, 3A, and 3B), but a thickness of ∼850 m has been measured in the Elko Hills (Haynes, 2003; Fig. 1B). There is no evidence from this study or from the interpretation of seismic reflection data in Huntington Valley that the Elko Formation and other pre-Miocene sedimentary rocks thicken significantly eastward into the west-dipping RMEH detachment fault (Satarugsa and Johnson, 2000; Haynes, 2003; Colgan et al., 2010). Interpretations based on the existing seismic lines (Satarugsa and Johnson, 2000) are challenging due to the data quality.
Southward-migrating volcanism reached northern Nevada ca. 42.6 Ma (Brooks et al., 1995), and coincided with the beginning of a Cenozoic phase of crustal melting (ca. 42–29 Ma) deep in the MCC infrastructure now exposed in the northern RMEH (e.g., Wright and Snoke, 1993; McGrew and Snee, 1994; MacCready et al., 1997; Howard et al., 2011; Drew, 2013; Fig. 10). Volcanism reached the mapped area ca. 38 Ma (Ressel and Henry, 2006; this study). A tuff dated as 37.9 ± 0.5 Ma (Table 2) near the top of the Elko Formation dates the end of deposition of the Elko Formation ca. 38 Ma. Because the ages of ash flow tuffs in the basal overlying volcanic package overlap this age within error, the uppermost dated tuff in the Elko Formation may have erupted from the Robinson Mountain volcanic field or from slightly to the north, near Emigrant Pass (Ressel and Henry, 2006; Figs. 1B and 3B; Table 2). No unconformity is observed between strata assigned to the Elko Formation and the overlying Eocene volcanic rocks of the Robinson Mountain volcanic field (Fig. 2), formerly mapped as the Indian Well Formation by Smith and Ketner (1978).
Ignimbrites of the Robinson Mountain volcanic field were erupted between ca. 38.5 and 37.3 Ma (Table 2) in a subaerial or shallow lacustrine environment, and are >1000 m thick in parts of the mapped area (Fig. 3E). The ignimbrites were likely erupted from vents in the Robinson Mountain area, as suggested by the presence of voluminous Eocene rhyolite domes of similar mineralogy and geochemistry (Supplemental File) and the lack of mapped calderas of this age (Henry and John, 2013). Eocene rocks of the Robinson Mountain volcanic field may have been deposited as far west as the eastern flanks of the present Cortez Mountains (Figs. 1B and 11), where patches of strata previously mapped as Indian Well Formation include basalt flows and white Paleogene lapilli tuffs exposed beneath Quaternary and late Cenozoic sedimentary cover (Smith and Ketner, 1976, 1978). The slightly younger ages of the crosscutting subvolcanic intrusions in the mapped area (ca. 37.6–36.8 Ma; Fig. 2; Table 2) document magmatism continuing to this time. Once erupted, the resistant volcanic rocks protected underlying sedimentary strata from subsequent erosion. Westward tilting (10°–15°) occurred after the eruption of 37.3 ± 0.3 Ma ignimbrites. Following tilting, Eocene volcanic rocks were covered with cross-bedded sandstone with a maximum depositional age of 33.9 ± 0.4 Ma (Figs. 2 and 11; Table 2).
Despite evidence that metamorphism, partial melting of the crust, and intrusion of magmas continued at depth in the developing RMEH until at least 29 Ma (e.g., Wright and Snoke, 1993; McGrew and Snee, 1994; MacCready et al., 1997; Howard et al., 2011; Fig. 10), volcanic activity at the surface ended ca. 36 Ma (Ressel and Henry, 2006; Walker et al., 2006; du Bray, 2007; this study; Fig. 10; Table 2). The thin tuff of Hackwood Ranch was erupted across the sediments overlying the Robinson Mountain volcanic rocks at 31.1 ± 0.3 Ma (Figs. 10 and 11), but, as discussed, it is likely that it was erupted from a distant caldera, probably near the Desatoya or Clan Alpine Mountains (Faulds et al., 2005; Henry et al., 2012; Henry and John, 2013). Based on these relationships, it is clear that magmatism at depth in the crust did not result in surface eruptions between 36 and 29 Ma. The tuff of Hackwood Ranch was subsequently tilted gently westward (Figs. 2 and 11). Because exposures of Eocene and Oligocene volcanic rocks are absent east of Cedar Ridge and Red Spring (Figs. 2 and 3A), they may have been deposited thinly or not at all (Fig. 11). If appreciable thicknesses of volcanic rocks had been deposited east of Cedar Ridge, they must have been eroded away before middle Miocene time.
The gently east dipping, mostly Miocene Humboldt Formation overlies the gently west dipping Oligocene tuff of Hackwood Ranch with ∼30° of angular discordance (Fig. 2). This angular discordance probably developed as a result of westward tilting between ca. 31.1 and 24.4 Ma. This early phase of westward tilting could have resulted from slip on east-dipping faults west of the exposures of the tuff of Hackwood Ranch, and/or from differential uplift of the region that was to become the RMEH MCC relative to surrounding areas due to crustal flow and/or diapiric rise of rocks at depth beneath what was to become the Ruby Mountains (Figs. 1B, 2, 10, and 11). Basin development along the west side of the RMEH (Fig. 1B) is heralded by deposition of the Humboldt Formation perhaps as early as 24.4 Ma. The basin filled mostly between ca. 16 and 12 Ma, and the youngest deposits are dated as 8.2 Ma (Colgan et al., 2010; this study). The youngest biotites previously dated within the RMEH, along the western side of the range, yielded early Miocene (ca. 24–21 Ma; Fig. 10) 40Ar/39Ar ages (McGrew and Snee, 1994). These ages are compatible with stratigraphic evidence for the onset of uplift by faulting in the Miocene and, because they may represent ages set within the partial retention zone for argon in biotite, they would represent maximum ages for faulting.
Although extensional faulting may have begun as early as 24.4 Ma, it accelerated ca. 16–15 Ma, as recorded by the onset of rapid deposition in the study area. This conclusion is in close agreement with recent findings, based on several techniques, that rapid uplift of the southern RMEH began ca. 17–15 Ma (Colgan et al., 2010). More than 2100 m of sediment were deposited across just the mapped area during Miocene time. The Humboldt Formation thickens eastward toward the RMEH (Satarugsa and Johnson, 2000; Fig. 1B), from where its younger strata are now known to have been sourced (Colgan et al., 2010; this study). An unroofing signature is recorded in the Humboldt Formation with the arrival ca. 14.2 Ma or later of metamorphic clasts and a detrital zircon age spectrum consistent with derivation of sediment from Proterozoic and early Paleozoic rocks now exposed in the southern RMEH.
Thus, prior to ca. 16 Ma, the supracrustal history of the sedimentary basin on the western flank of the RMEH MCC provides evidence for only minor tilting near and within the Piñon Range, or the western flank of the future RMEH MCC. It is possible that these tilts are a result of east-dipping normal faults, or they might reflect adjustments in the upper crust related to Eocene and Oligocene flow at depth, and thus could represent the western side of a dome-like uplift.
Our results, which indicate only minor surface deformation in the time span from the Late Cretaceous to the Miocene, help to underscore the dilemma posed by the suggestion that large-magnitude thrusting followed by large-magnitude extension occurred across the region of the RMEH in the Late Cretaceous to early Cenozoic (e.g., Coney and Harms, 1984; Vandervoort and Schmitt, 1990; Hodges and Walker, 1992; Mueller and Snoke, 1993; Camilleri, 1996; Camilleri and Chamberlain, 1997; McGrew et al., 2000; Howard, 2003; Hallett and Spear, 2014, 2015). In particular, the growing detrital zircon record obtained from sandstones and conglomeratic units at the base of the Cenozoic unconformity across the northern BRP does not show evidence for recycling and thus uplift of the lower part of the shelf succession (late Precambrian to Cambrian) or widespread deposition related to high-angle faulting and half-graben formation prior to Eocene time (e.g., Druschke et al., 2009, 2011; Konstantinou et al., 2012; Ruksznis, 2015; this study). These findings in turn challenge proposals that large-magnitude extension in the hinterland of the Mesozoic Sevier fold and thrust belt was driven by the onset of melting and extensional collapse of structurally overthickened, gravitationally unstable crust (e.g., Molnar and Chen, 1983; Coney and Harms, 1984; Hodges and Walker, 1992). Instead, major extension near the RMEH and Piñon Range, as documented by the deposits studied here, probably began in middle Miocene time, more than 50 m.y. after Mesozoic crustal thickening, metamorphism, and magmatism peaked in the Sevier hinterland (Miller and Gans, 1989). Even allowing 20 m.y. for the relaxation of isotherms in thickened crust (e.g., Camilleri and Chamberlain, 1997; DeCelles, 2004; Mattinson et al., 2007), Miocene extension began more than 30 m.y. after attainment of these elevated temperatures.
In contrast to the view from the upper crust discussed here, pressure-temperature-time paths determined using metamorphic mineral assemblages and U-Pb monazite cooling ages yield evidence for uplift and cooling of originally deep-seated northern RMEH rocks beginning during the Late Cretaceous or even earlier (Dallmeyer et al., 1986; Hodges et al., 1992; Hallett and Spear, 2014, 2015; Fig. 10). Fission-track, U-Pb, and 40Ar/39Ar thermochronology data suggest that uplift of these rocks continued through Oligocene time (e.g., Dallmeyer et al., 1986; Dokka et al., 1986; Wright and Snoke, 1993; McGrew and Snee, 1994; MacCready et al., 1997; McGrew et al., 2000; Snoke et al., 2004; Premo et al., 2005; Fig. 10). Further work is needed to understand the nature and location of structures that accommodated uplift of these rocks and to reconcile these structures with what is now known about the surface geology and topography of the area near the RMEH during this time span. Our data from supracrustal rocks do not preclude uplift of the deeper parts of the crust by a process such as diapirism and/or crustal thinning, as modeled by Rey et al. (2009), who suggested that mechanical decoupling occurred between the brittle upper crust and the ductile crust below (see also MacCready et al., 1997). This could explain the relatively low degrees of brittle faulting and topographic changes prior to Miocene extension.
Implications for Estimates of Paleotopography, Paleoelevation, and Paleoclimate
Stable isotope studies that address the evolution of Paleocene to Holocene climate, elevation, and topographic relief across the Sevier hinterland utilize and are heavily influenced by data collected from the Elko Basin (Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Feng et al., 2013; Mulch et al., 2015). These studies argue for a rapid elevation gain that swept southward through the latitude of the Elko Basin in the Eocene, tracking the southward migration of volcanism following the end of flat-slab subduction (Gans, 1990; Mix et al., 2011; Chamberlain et al., 2012). Chamberlain et al. (2012) and Mulch et al. (2015) suggested that the shift in stable isotope values documented in the Elko Basin for ca. 40–38 Ma also heralds the formation of rugged topography with peaks >4000 m across the region that affected climate and rainfall.
Our findings do not bear on absolute elevations, but call into question the suggestion that rugged topographic relief with peaks >4000 m formed here in the Eocene (Chamberlain et al., 2012; Mulch et al., 2015). Instead, geologic data suggest that the region had relatively low relief and was covered by volcanic rocks that underwent low magnitudes of erosion after volcanism ceased (Henry, 2008; Best et al., 2009; Van Buer et al., 2009; Henry et al., 2012; Konstantinou et al., 2012; Long, 2012; this study). A major implication of our remapping and more detailed dating of the section is that the measured isotopic shift reported by Horton et al. (2004), Mix et al. (2011), Chamberlain et al. (2012), and Mulch et al. (2015) is no longer constrained to have occurred in the Eocene, but could have taken place anytime between ca. 40 and 16 Ma. Thus the isotopic data (Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Mulch et al., 2015), when tied to the revised stratigraphy presented here, are in general agreement with a low-relief Eocene volcanic tableland, followed much later in the Miocene by the development of higher mean elevations and mountain ranges (rugged topography) during Basin and Range faulting (e.g., Miller et al., 1999; Colgan et al., 2010; Konstantinou et al., 2012; this study).
Our detailed geological study of Huntington Valley and the eastern Piñon Range, together with supporting geochronology, reveals that the upper crust and supracrustal section exposed in this area experienced mostly tectonic quiescence from Late Cretaceous through Oligocene time. It is striking that this interval of surficial tectonic quiescence coincides with partial melting, metamorphism, and lower crustal flow at depth in the developing RMEH MCC (e.g., Wright and Snoke, 1993; McGrew and Snee, 1994; MacCready et al., 1997; Howard et al., 2011; Figs. 10 and 11).
Deposition was minor or absent in central Huntington Valley and the eastern Piñon Range during the time span of Mesozoic crustal shortening. Regional erosion followed, exposing gently folded late Paleozoic and Mesozoic (Triassic) strata that are now beneath the Tertiary unconformity. Redbeds and limestone were deposited unconformably and discontinuously over the older folded strata. The detrital zircon signatures of the redbeds and those of the overlying Eocene Elko Formation suggest recycling of zircons specifically from Permian and Triassic miogeoclinal strata. The Eocene Elko Formation is only ∼180 m thick at its greatest in the mapped area, but is as much as ∼850 m regionally (Haynes, 2003). The Elko Formation contains ca. 46–38 Ma detrital zircon populations that represent air fall derived from volcanic eruptions north of the Snake River Plain (46–42 Ma) and locally (39–38 Ma). Local eruption of rhyodacite ignimbrites, rhyolite domes and flows, and intrusion of subvolcanic rhyolite of the ca. 38.5–36.8 Ma Robinson Mountain volcanic field followed deposition of the Elko Formation.
Westward tilting of 10°–15° occurred before deposition of thin 33.9 Ma (or younger) sediments and eruption of the 31.08 ± 0.47 Ma rhyolitic tuff of Hackwood Ranch, which was probably derived from a distant caldera. This was followed by additional, apparently localized, westward tilting by ≥10°–15° that occurred before ca. 24.4 Ma. At least some of the westward tilting was probably caused by normal slip on east-dipping faults west of the study area, but it is also possible that tilting could be related to uplift of a developing RMEH gneiss dome at greater depth in the crust. If so, our data may suggest decoupling of events in the deep and shallow crust and may allow for earlier (diapiric) rise of metamorphic rocks, such as is now documented for the Albion–Raft River–Grouse Creek Mountains MCC (Strickland et al., 2011; Konstantinou et al., 2012, 2013). Rapid slip began ca. 16 Ma on a west-dipping normal fault bounding the west side of the RMEH (Colgan et al., 2010); this led to deposition of the Humboldt Formation from ca. 15.8 to 12 Ma, with minor deposition beginning perhaps as early as 24 Ma and continuing until ca. 8.2 Ma or later. More than 2000 m of Miocene Humboldt Formation strata were deposited across the study area, and deposits of this basin thicken eastward toward the RMEH. Detrital zircon populations and metamorphic clasts thought to demonstrate RMEH provenance are first detected in 14.2 Ma (or younger) Humboldt Formation deposits, which constrain the timing of RMEH unroofing. Our conclusions parallel earlier discoveries that the metamorphic rocks in MCCs of the northern BRP were uplifted in their final stage only in Miocene time (Miller et al., 1999; Colgan et al., 2010; Konstantinou et al., 2012, 2013; Ruksznis, 2015).
Our work also has implications for the paleogeographic evolution of the hinterland of the Sevier fold and thrust belt. We reassign a thick succession of sedimentary strata previously assigned to the Eocene–Oligocene Indian Well Formation to the lower part of the mostly Miocene Humboldt Formation on the basis of geologic mapping and extensive geochronologic constraints. The reassignment of these strata compels a reinterpretation of stable isotope–based paleoelevation estimates and raises the prospect that the proposed large and rapid elevation gain (∼2.5 km in <2 m.y.) in the Eocene (e.g., Chamberlain et al., 2012) could have taken place over any time interval between the middle Eocene and middle Miocene. Although we cannot establish paleoelevations of this region based on our current work, the stratigraphic details of the Cretaceous(?) to Miocene section studied precludes the conclusion arising from stable isotope studies that rugged topography was established across this region in the Eocene (Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Feng et al., 2013; Mulch et al., 2015). Instead, the region was topographically subdued and covered with volcanic rocks until the Miocene, when rugged topography first formed.
This research was supported by Stanford University, U.S. Geological Survey EdMap award G12AC20189 to Lund Snee and Miller, and National Science Foundation Tectonics Program grant 1322084 to Miller. We thank Alan J. McGrew and an anonymous reviewer for constructive reviews, and Terry Pavlis for thoughtful and responsive editorial assistance. Our research benefitted from discussions with J.P. Colgan, D.A. John, C.D. Henry, T.A. Dumitru, and G.A. Thompson. We thank H. Dudley, M. Erviti, J. Knox, M. Raftrey, B. Girma, S. Xiao, and the students of Stanford Field Class GES 190 for their able assistance.