Neogene (Miocene–Pliocene) sedimentary rocks of the northeastern Sierra Nevada were deposited in small basins that formed in response to volcanic and tectonic activity along the eastern margin of the Sierra. These strata record an early phase (ca. 11–10 Ma) of extension and rapid sedimentation of boulder conglomerates and debrites deposited on alluvial fans, followed by fluvio-lacustrine sedimentation and nearby volcanic arc activity but tectonic quiescence, until ∼ 2.6 Ma. The fossil record in these rocks documents a warmer, wetter climate featuring large mammals and lacking the Sierran orographic rain shadow that dominates climate today on the eastern edge of the Sierra. This record of a general lack of paleo-relief across the eastern margin of the Sierra Nevada is consistent with evidence presented elsewhere that there was not a significant topographic barrier between the Pacific Ocean and the interior of the continent east of the Sierra before ∼ 2.6 Ma. However, these sediments do not record an integrated drainage system either to the east into the Great Basin like the modern Truckee River, or to the west across the Sierra like the ancestral Feather and Yuba rivers. The Neogene Reno-Verdi basin was one of several, scattered endorheic (i.e., internally drained) basins occupying this part of the Cascade intra-arc and back-arc area.


The details of the Cenozoic evolution of the Sierra Nevada mountain range, and of the tectonic mechanisms responsible for it, have long been subjects of interest and controversy. Tectonic conditions in the Sierra have changed as the North American-Farallon-Pacific plate boundary to the west evolved from convergent to transform (e.g., Atwater, 1970; Atwater and Stock, 1998). Added complications along the eastern margin of the Sierra block are the development of the Basin and Range Province (e.g., Surpless et al., 2002) and the formation of the Walker Lane as a complex tectonic boundary in its own right (e.g., Atwater and Stock, 1998; Faulds et al., 2005a, 2005b; Faulds and Henry, 2006; Busby et al., 2010). These tectonic controls have varied in space, primarily N-S along the range, as well as in time. Distinguishing their effects to better understand the significance of each will require careful paleogeographic and structural studies, as well as good age control, along the length of the range. In this study, we focus on the eastern margin of the northern Sierra, along the Nevada-California state line (Fig. 1).

The Neogene sedimentary basins of northwestern Nevada and northeastern California provide important constraints on the evolution of the northern Sierra Nevada. Geographically, they occur across the eastern boundary of the Sierra: There are Neogene sedimentary rocks within the northern Sierra Nevada (e.g., Mass et al., 2009), in the Walker Lane (e.g., Schwartz, 2001; Trexler et al., 2009), and in a faulted region between the Walker Lane and the Sierran frontal fault system (e.g., Gilbert and Reynolds, 1973; Golia and Stewart, 1984; Cashman et al., 2009). The Neogene sedimentary rocks were deposited during the time period when both the Walker Lane and the Sierran frontal fault system were becoming active at this latitude. The sedimentary rocks preserve evidence of the evolution of these features by recording changes in depositional settings and paleogeography with time.

In this paper, we present the geologic record from an endorheic Neogene sedimentary basin including the area now occupied by Reno and Verdi, Nevada. After summarizing the tectonic setting and previous work, we describe the internal stratigraphy of the Neogene Verdi basin. We document the sedimentology, adding the fossil record and age control where available. We synthesize the accumulating evidence for the late Miocene–Pliocene topography, climate, ecology, and depositional settings that dominated this part of what is now the northeastern Sierra Nevada at the latitude of Reno, Nevada. We show that this basin history entirely predates the modern relief of the adjacent Carson Range and the nearby northern Sierra Nevada relative to that of Reno and the Truckee Meadows. We therefore constrain formation of the dramatic topographic relief in the northeastern Sierra Nevada to be entirely younger than 2.6 Ma, the age of the youngest dated Neogene strata in the basin. This paper is a synthesis of the work of the authors and many others, over the course of more than 100 years of geologic study in the area.


It is important to consider the formation of the Sierra Nevada as a high mountain range separately from the creation of structural relief along the Sierran frontal fault system, the eastern margin of the range. In addition, it is probably an oversimplification to consider either of these to be a single event. Many studies have used different methods to determine the timing, rate, or history of uplift of the Sierra Nevada, resulting in complex and sometimes contradictory interpretations. We restrict the following summary and discussion to the northern Sierra Nevada.

Advocates of Paleocene (possibly as old as Cretaceous) uplift of the Sierra Nevada cite several different lines of evidence: (U-Th) He thermochronology along a 100 km transect perpendicular to the northern Sierra Nevada has been interpreted to record rapid uplift 60–90 Ma ago, followed by a long period of slower exhumation from early Paleogene to the present (Cecil et al., 2006). These authors point out that their result is consistent with a thick, widespread, and well-documented lateritic paleosol that is developed throughout the northern Sierra Nevada. The paleosol is found at the base of Eocene sedimentary strata, so it records pre-Eocene exposure and weathering. Stable isotope ratios of hydrogen in kaolinite of Eocene ancestral Yuba River sediments record the effect of topography on precipitation; they indicate a high Sierra Nevada 40–50 Ma ago (Mulch et al., 2006). The sedimentology of coarse Eocene fluvial strata on the northwest Sierra slope shows that formation of these rocks required uplift and locally steep gradients (Cecil et al., 2004; Cecil et al., 2010; Cassel and Graham, 2011). Paleobotanical studies of leaf morphology suggest a high Sierra Nevada in the Miocene that subsequently collapsed (Wolfe, 1994; Wolfe et al., 1997). More recent studies based on stomatal density in middle Miocene leaves indicate that the Sierra Nevada was already high before Miocene time (Kouwenberg et al., 2007). Paleoaltimetry based on studies of stable isotopes of hydrogen in volcanic glass in ignimbrites is interpreted as indicating a high and steep Sierra Nevada 28–31 Ma ago (Cassel et al., 2009b). However, the isotopic shift attributed to an early orographic effect might be equally due to a change in climate from a wetter and warmer Paleogene changing to a cooler dryer Neogene (Molnar, 2010).

In apparent contradiction to arguments for a high Paleogene Sierra, an Eocene–Oligocene fluvial system draining westward from central Nevada across the longitude of the present Sierra has long been cited as evidence that the Sierra Nevada was not a topographic barrier to the Great Basin in Paleogene time (Fig. 1) (e.g., Durrell, 1957, 1959, 1966; Bateman and Wahrhaftig, 1966; see the thorough summary in Garside et al., 2005, and references therein). Evidence for an origin of these west-flowing rivers east of the Sierra crest includes the size of the river systems and the presence of clasts that did not originate in the Sierra, including some clasts derived from Nevada. These gravels, known as the deposits of the Tertiary (or “ancestral”) Yuba, American, and Feather rivers, contain auriferous gravels that were hydraulically mined in the nineteenth century, so their extent and outcrop pattern are well documented. Today these conglomerates are elevated relative to their modern fluvial counterparts, supporting the interpretation that some Sierran uplift postdated the ancestral, through-going river systems. Cassel et al. (2010) have been able to identify detrital zircons in these rocks that they attribute to sources in central Nevada, and they invoke a high interior craton that drained from central Nevada across the Sierra Nevada to the Pacific Ocean. Note that one of these paleovalleys has been projected to flow directly across the Reno-Verdi basin study area (Fig. 1).

The “Nevadaplano” hypothesis, the suggestion that an early, high Sierra Nevada could have extended east into central and eastern Nevada (first proposed by DeCelles, 2004, and Robinson and McQuarrie, 2004, and reviewed by Ernst, 2010), can accommodate both a relatively high Paleogene Sierra and through-going Paleogene fluvial systems. Convincing evidence of this geography includes west-draining paleovalleys that cross the range, incised into the granite and country rocks. These are filled with ignimbrites and Eocene–Oligocene volcanogenic sedimentary rocks (e.g., Garside et al., 2005; Busby and Putirka, 2009; Cassel and Graham, 2009; Cassel et al., 2009a; Egger et al., 2009; Gorny et al., 2009; Hagan et al., 2009; Henry et al., 2009; Koener and Busby, 2009; Schweickert, 2009; Busby et al., 2010; Henry and Faulds, 2010).

The Oligocene western Nevadaplano, crossed by valleys running to the Pacific Ocean, was buried in the early Miocene by andesitic lavas and volcaniclastic deposits. Through-going drainage to the Pacific Ocean was cut off by the return of the volcanic arc to the longitude of the modern Sierra Nevada. In middle Miocene time, a short episode of local normal faulting and mafic volcanic activity (e.g., Henry and Perkins, 2001) disrupted the landscape, and in many areas small- to moderate-size sedimentary basins filled in topographic lows (Fig. 1, Neogene basins in green). These Neogene sediments rest on a variety of rocks and on a topographically uneven surface; the Miocene volcanic section ranges from hundreds of meters thick to locally missing below Neogene basin strata.

The Miocene–Pliocene sedimentary deposits of the Verdi basin predate the modern relief at the northeastern boundary of the Sierra Nevada. Pliocene to Holocene sedimentary rocks and structures document the formation of this topography. See the companion paper (Cashman et al., 2012) for a discussion of this more recent geologic history.


The Neogene sedimentary rocks of northwestern Nevada-northeastern California, including but not limited to the green areas on Figure 1, were collectively called the “Truckee Formation” by the 40th Parallel Survey (King, 1878). We have shown that although strata of these areas are in part correlative, they represent different localized basins, each with its own history (Trexler et al., 2000). The internal stratigraphy of the Verdi basin was first described in a report of an oil exploration well in western Reno, which drilled through 1890 ft. (576 m) of Neogene section (Campbell, 1908). U.S. Geological Survey 15′ quadrangle mapping 50 years later established the distribution of the Neogene sedimentary rocks on the flanks of the Carson Range (Thompson and White, 1964). A generalized stratigraphy of these rocks near Reno informally named them “the Sandstone of Hunter Creek” (Bingler, 1965). Mostly unresistant, they crop out in hillsides, gully walls, strath terrace risers, road cuts, railroad cuts, and irrigation ditches throughout the valleys of western Reno, Mogul, and Verdi in northwest Nevada (Fig. 2). Bingler’s stratigraphy—lower and upper clastic units separated by a middle diatomaceous section—was used in mapping the Verdi 7.5′ quadrangle (Bell and Garside, 1987). Our subsequent work (e.g., Trexler et al., 2000) has shown that a uniform vertical stratigraphy does not apply everywhere, but age correlation and synthesis of sections throughout the basin provide a detailed history. We retain the informal term “Sandstone of Hunter Creek” for these rocks because it has been applied only to the rocks exposed in and west of Reno, and it does not imply correlation or continuity with other Neogene rocks in the region.

The Neogene sedimentary rocks of the Reno-Verdi basin are highly varied in lithofacies and thickness (Fig. 3). We have synthesized internal stratigraphy using age control where possible. It is important to note that while there are some very general stratigraphic trends, as noted by earlier workers (e.g., Campbell, 1908), lithofacies cannot reliably be used to correlate these stratigraphies in detail.

Coarse-Grained, Basal Section

The older strata of the Neogene Verdi basin are characterized by coarse-grained sedimentary rocks and are locally interbedded with volcanic rocks (Fig. 3). They are exposed around the foothills of the Truckee River corridor west of Reno, and in uplifted fault blocks between Verdi and Reno (Fig. 2). The basal section is coarse-grained everywhere it is exposed, but its thickness is inconsistent. Coarse conglomerate dominates many basal intervals, and is interbedded with sandstone, siltstone, and diatomaceous sediments. Debrites (debris flow deposits, see Gani, 2004) are also common in the basal section. The basal section depositionally overlies, and in places is also interbedded with, the youngest andesitic volcanic rocks in the area, the Kate Peak and Alta formations (Bell and Garside, 1987). In several places, the basal section contains basaltic andesite flows. In addition, the lower part of the section in the north Carson Range foothills contains at least one and possibly several syneruptive lahar deposits, indicating local active intermediate volcanism early in the basin history (best exposed at Mayberry Park, Figure 2, locality C). The youngest lahar is andesitic and monolithologic. It exhibits inverse grading and contains several amalgamated flow units that are tens of meters thick. Radially cracked clasts are evidence that these debrites represent a hot lahar associated with an eruption, rather than a later mudflow that incorporated existing volcanic material.

The basal conglomeratic interval is best studied in the southwest part of the preserved basin strata, where it is relatively well exposed and over 100 m thick. Paleocurrent directions, determined from clast imbrication, indicate that drainage flowed east (Fig. 4A). Conglomerate clasts are well rounded and are commonly as large as 20–30 cm. Cobbles are predominantly intermediate volcanic rocks. Elsewhere around the margin of basin exposure, basal conglomerates are relatively thin, less than 10 m thick. Exposures of these conglomerates are generally very poor because they are only weakly cemented, and ravel easily. However, their presence can be mapped with confidence from the abundant well-rounded cobbles at the ground surface.

An anomalous basal section is exposed in horst blocks in the center of the Truckee River corridor (Fig. 2), where conglomerate directly overlies granite and granite regolith. Here, the basal section contains abundant granite boulders as well as volcanic clasts and rare metavolcanic cobbles. This is unusual for Verdi basin sediments, which in most places lacked a source of granitic clasts. The granite basement here is deeply weathered beneath the Miocene sediments, with spheroidally weathered blocks still in place and rooted to basement (Fig. 5). This weathering pattern suggests a very long period of erosion and exposure before the basal Miocene conglomerate was deposited. Excellent road-cut exposures display both monolithologic granite debris-flow and coarse heterolithic fluvial or alluvial fan conglomerate beds in an interval 30–50 m thick (Fig. 6).

Several lines of evidence illustrate that the basal contact of the Neogene section is a buttress unconformity. The entire arc-volcanic section, except a nearby basaltic andesite flow, is missing at the nonconformity on the horst block. However, Kate Peak and Alta formation volcanics are quite thick less than 3 km to the south and north, respectively. The lowest Sandstone of Hunter Creek invariably contains conglomerate derived from the Kate Peak or Alta volcanics. The age of the base of the Neogene clastic section is younger from west to east along the preserved margins, although bedding attitudes are conformable. In the center of the exposed basin, at the nonconformity with granite basement, the age is ca. 10 Ma (C.D. Henry, reported in Trexler et al., 2000). To the east it is younger, projected to be as young as 4–5 Ma (based on correlation with dated sections, discussed below) where Neogene exposures disappear under the glacial outwash that fills the Reno basin. Therefore, Miocene volcanic topography consisted of volcanic strata unevenly deposited around eruptive centers. The Neogene Sandstone of Hunter Creek originally buried much of this volcanic topography. Younger parts of the Hunter Creek section successively buried topographically higher areas. All is now being exhumed in the modern Truckee River canyon and on the flanks of the surrounding ranges.

The basal section of the Hunter Creek Sandstone is anomalously thin and fine-grained along the north and west margins of the Neogene basin exposures. Just to the northwest of the horst-block area described above, basal Miocene sediments are conglomerate and sandstone, overlain by diatomite. Here, basaltic andesite flows (10.3 Ma) (C.D. Henry, reported in Trexler et al., 2000) are intercalated with diatomite beds locally, and have baked them to a pink color and porcelainous texture. The basalt is scoriaceous, having apparently erupted into a lake. These diatomite beds are among the oldest lacustrine deposits in the Verdi basin, and occur primarily in the northwestern part of the preserved basin.

Fossil material collected from the lower part of the Verdi basin section identified by Kelly (2001) and Kelly and Secord (2009) provides insight into the Miocene paleoenvironment. It includes Dinohippus cf. D. leardi (horse), Camelidae (camel), ?Antilocapridae (related to modern antelope and giraffe), and Mammutidae or Gomphotheriidae (ancestors of elephants). These animals include both browsers and grazers, and their presence suggests a stable grassland and open woodland environment with riparian vegetation and thicker woodlands along perennial streams and lakes (Kelly and Secord, 2009).

The shift from predominantly volcanic and volcaniclastic strata of the Kate Peak Formation to the dominantly sedimentary Sandstone of Hunter Creek appears to record a cessation of local volcanism. The Kate Peak Formation contains abundant sedimentary rocks that were deposited in a variety of volcano-dominated environments throughout the upper Miocene (e.g., Mass, 2005; Cousens et al., 2008; Mass et al., 2009), and the transition upward includes andesitic lahars, basaltic andesite flows, and pumiceous and scoriaceous sandstone. We interpret this transition as a shift in local volcanic centers. The change in magmatic composition from arc volcanism to more mafic flows may have accompanied an extensional tectonic episode ca. 10–10.5 Ma (Henry and Perkins, 2001, and discussed below).

The lower interval of the Sandstone of Hunter Creek is well constrained in age because of the relatively abundant intercalated volcanic rocks. Isotopic age determinations by C. Henry (summarized in Trexler et al. [2000] and Henry and Perkins [2001]) include the following (see Figures 2 and 3 for sample locations):

(1) The youngest andesitic volcanic rocks in the area, at and within the base of the Neogene Verdi basin section, which are at least as young as 12.1 Ma below the Sandstone of Hunter Creek and include rocks dated as 11.72 ± 0.11 or 11.39 ± 0.11 Ma within the section.

(2) A thin basaltic andesite flow found stratigraphically just above the lahar at the base of the section, which dates volcanic activity here at no younger than 10.2 Ma.

(3) Basaltic andesite flows that crop out near the horst blocks in the central area appear to be interbedded low in the stratigraphic section and are dated at 10.3 Ma.

(4) In several places, the basal section contains basaltic andesite flows that range from 10.4 to 10.1 Ma.

Biostratigraphic age control on the mammal fossils collected here is not precise, but is consistent with isotopic determinations. Ages of this material are consistently middle to late Clarendonian (10–7 Ma) (Kelly and Secord, 2009).

Between 12 and 10.5 Ma, basin deposition began on a moderately high-relief basement surface of deeply weathered Mesozoic granite and fresh Cenozoic intermediate volcanic rocks. In the case of the latter, this depositional shift was a continuation of volcaniclastic and volcanic deposition. Clast imbrication of basal conglomerates of the Sandstone of Hunter Creek record a well-developed, east-directed river system (Fig. 4A) in a drainage catchment dominated by intermediate volcanic rocks, and notably lacking in exposed granite. Andesitic lahars and basaltic andesite flows (the latter ranging from 10.4 to 10.1 Ma) are locally intercalated with basal conglomerates, and small (<1 km diameter) lakes with local fan-deltas formed mainly in the northern and western part of the basin. Coarse, 3–5-m-thick granite-clast debris flows are intercalated with the basal conglomerates in the center of the Truckee River corridor, and record local uplift of granitic basement, probably along faults that were active at the time.

We interpret the mid-Miocene landscape as one of moderate local topographic relief with alluvial fans and some rock avalanches along the margins of fault-bounded valleys, with perennial lakes and energetic rivers, grasslands, and forested uplands. There was enough precipitation and vegetation to support a prolific mammal megafauna. The fluvial system flowed east, and the landscape must have been dominated by the active volcanic arc to the west, with nearby stratavolcanoes.

Middle Medium- to Fine-Grained Section

Above the basal conglomerates, the Hunter Creek section grades upward into interbedded pebbly litharenite, litharenite, siltstone, and muddy diatomite. In general, the coarser-grained rocks are thicker and dominate exposures in the southern part of the basin (roughly, south of the Truckee River) and finer-grained rocks dominate the northern part. The most continuous exposures of this part of the section are along the Steamboat irrigation ditch on the north flank of the Carson Range, and locally in railroad cuts south of the Truckee River and west of Mogul (Fig. 2).

Sedimentary fabric in coarser sandy beds includes abundant lower-flow regime features commensurate with grain size: ripple cross-lamination, trough cross-lamination, and small channel structures. Finer beds are mostly silty, diatomaceous in places, contain rooted horizons, and locally preserve abundant plant material. Tephras, scoria-pebble beds, and pumice-pebble beds in the sandstones and pebble conglomerates indicate that volcanic sources were still active nearby, but no volcanic flow-rocks occur in this part of the basin section.

Paleocurrent directions in this part of the section are mixed, but generally indicate drainage to the northwest, north, and northeast (Fig. 4B). Locally these indicators (most are bar foresets) are highly divergent, as would be expected for meandering fluvial systems and shifting delta lobes. Overall, however, they are consistent with streams draining, and deltas building, northward. Finer-grained rocks, including lacustrine diatomite, in northern exposures record the downstream termination of these drainage systems.

Mammal fossils provide direct age control from this part of the section, as well as independent paleoenvironmental information (Kelly, 2001; Kelly and Secord, 2009). Three relatively widespread localities within a few tens of meters stratigraphically above the basal conglomerates yield late Clarendonian fossil fragments (including a partial tusk of a gomphothere), and 10 additional localities higher in the fluvial/deltaic section yield fossil fragments from throughout the Hemphillian, including at least two species of camel and a horse (Kelly, 2001; Kelly and Secord, 2009) (see Figure 3 for North American Land-Mammal Ages [NALMA] time scale). Both browsing and grazing taxa are found throughout both the Clarendonian and Hemphillian portions of the section, indicating a fairly stable environment (Kelly, 2001; Kelly and Secord, 2009) that lasted for at least 5 million years.

Paleobotanical information fills out the depositional environmental description. The well-known “Verdi flora” (University of California plant locality 102) occurs in this part of the section, in a railroad cut on the south bank of the Truckee River (Figure 2, locality A; Axelrod, 1956, 1958). The Verdi flora comprises almost completely riparian and lake-border species. It is interpreted to be Hemphillian, based on paleobotanical correlation, supplemented by a K/Ar date of 5.85 Ma (recalculated) (Axelrod, 1958; Evernden et al., 1964; Schorn, 1994). These floras resemble those that are found on the edge of the yellow pine forest on the Sierra Nevada west slope today, with forests of pine and fir, riparian areas dominated by willow, and open savanna with local chaparral. Axelrod (1958) interpreted this environment as sub-humid with yearly rainfall from 18 to 20 in. (two to three times today’s rate) and up to 25 in. at the forest margin. The January temperature was 10 °F higher than it is today and the frost-free season at least two months longer (Axelrod, 1958). The paleoelevation estimate is ∼2500 ft. (760 m) based on comparative flora today on the west side of the Sierra (Axelrod, 1958). The modern elevation of these rocks is 4900 ft. (1490 m). In summary, the climate and elevation suggest to us that there was no orographic rain-shadow effect of a high Sierra Nevada in the Miocene–Pliocene.

Isotopic age control has not been established in the middle part of the Neogene section because necessary rock types are lacking. Age control comes from tephrochronology by M. Perkins (Trexler et al., 2000) of a tephra near the top of the middle section, below the upper thick diatomite section (Fig. 3). This tephra matches two chemically indistinguishable tephra of 4.5 and 4.8 Ma, so it is thought to be one of these two ages. Biostratigraphy from mammalian fossils in the middle section indicates a Hemphillian age (Kelly and Secord, 2009), which fits well with the tephra age determination.

The sandstone petrology and conglomerate composition of the Neogene Verdi basin rocks document their local provenance, and reveal the extent to which the Sierra Nevada crystalline core was exposed during Miocene and Pliocene time. Loose sand and grain-mount thin sections were examined quantitatively to determine provenance. These results are from a study by Queen (2008). The sand grains are >80% volcanic lithic grains, with low percentages of quartz and feldspar (Fig. 7). These results are uniform for all samples examined throughout the section. Clast counts of conglomerate in Neogene units reveal an abundance of volcanic clasts ranging from andesite to basalt. Metamorphic rocks are rare; the source of these metavolcanic rocks is most likely Mesozoic roof-pendants on Peavine Mountain to the north, or in the nearby Sierra. Granitoid crystalline clasts do not occur in Neogene conglomerates, except at the very base of the section.

These petrologic data demonstrate that except for a brief time at the beginning of deposition, and only very locally, crystalline Sierran basement was not a significant sediment source for the Miocene–Pliocene Verdi basin. This finding contrasts with the Eocene–Oligocene sediments shed westward from the Sierran core into Eocene paleovalleys (discussed above), which are composed mostly of crystalline-source debris. However, these sedimentary rocks are older than the Verdi basin deposits and are not preserved in the Verdi area; they predate the onset of intermediate arc volcanism near the Sierran crest. We also note that the granite basement exposed today in the Verdi basin area and nearby Carson Range is deeply weathered at its contact with the overlying Neogene sedimentary-cover contact. Granite subsequently exposed in the Pleistocene is much less weathered.

Taken together, we interpret this provenance history as a long period of Paleogene uplift and stripping of cover rocks, exposing and deeply weathering the granite and metamorphic basement. In the early Neogene, arc volcanism buried much of this surface, but buried it unevenly, primarily around volcanic centers. Late Miocene and Pliocene basin sediments continued this burial among and around the volcanic edifices, resulting in the locally isolated but regionally coeval Neogene basin record we see today (e.g., Trexler et al., 2000; Oldow and Cashman, 2009).

The rocks of the middle part of the Hunter Creek section are consistent with an internally drained fluvial and deltaic-lacustrine system, with sources of sediment generally to the south (Fig. 4B). The finer-grained and more diatomaceous rocks near the northern edge of the basin (and also near the northwestern edge, probably for only a short time) record a shallow lake or lakes at the termination of these drainage systems. Fluvial intervals have a high sand-to-mud ratio, and sedimentary structures indicate meandering rivers a meter or two deep with unstable banks and frequent channel avulsion. Deltaic intervals contain sandy units with foreset structures consistent with small, Gilbert-type deltas, interbedded with diatomaceous silts representing interlobe swamps and floodplains. Two lines of evidence suggest local and active volcanic sources of sediment: Tephras are coarse and crystal-rich, and conglomerate contains relatively fragile scoria and pumice pebbles. No volcanic rocks occur in this part of the basin section. The early, western, basin lakes were filled by delta and floodplain deposits. By 5 Ma, sedimentation throughout the basin was dominated by sandy, meandering fluvial systems with associated lakes in the northern part of the basin (Fig. 8). Sediments younger than this are not preserved in the western part of the basin (i.e., west of Mogul, Fig. 2). The preserved sedimentary record shifts to the eastern part of the basin, where the section from 4.5 to < 2.6 Ma is preserved.

Upper Diatomaceous Section

The upper part of the stratigraphic section is dominated by diatomite, siltstone, and sandstone. Road cuts along old U.S. 40 (now 4th street, the locality is known as “Chalk Bluffs” on the Verdi 7.5′ topographic quadrangle, locality B on Figure 2) expose a 190-m-thick continuous section of diatomite (Figs. 2 and 3). The diatomite preserves many thin, crystal-lithic and glassy tephras. Tephra beds preserve diagnostic features of air-fall debris beds: lithic-ash couplets, graded grain-size, symmetric loading, and fluid injection features with no evidence for slope failure or lateral transport.

Up-section east of the 4th Street road cut, exposure is poor but scattered outcrops preserve beds of litharenite, siltstone, and silty diatomite. This interval contains small-scale, lower-flow regime current structures and common soft-sediment deformation features including ball-and-pillow structures and soft-sediment folds, due to fluid escape. Finally, near the eastern limit of exposure, the section ends in an erosional unconformity discussed in detail in a companion paper by Cashman et al. (2012).

To the south, in the northeast Carson Range foothills (southwest of Reno), correlative strata consist of diatomite intervals tens of meters thick interbedded with meter-thick litharenite beds. These sandy intervals preserve tabular foreset and trough cross-bedding, and are commonly deformed by soft-sediment, fluid-escape features. The section becomes dominated by sand upward, where it is buried by Quaternary glacial outwash deposits or is covered by modern construction.

Age control in the upper section comes from crystal-lithic tephras and mammal fossils (Trexler et al., 2000). One tephrochronology sample is from the sandy section low in the diatomite interval, and is 4.5 or 4.8 Ma (discussed above). Two tephras fairly high in the diatomite section are 3.16 and 3.06 Ma, respectively (Trexler et al., 2000). Using these tephras and measured sections, we have calculated the accumulation rate for the diatomite interval containing these tephras at ∼270 m/m.y. While this accumulation rate is fast, sedimentology shows no evidence for active, basin-bounding faults.

New 40Ar/39Ar analyses of mineral separates from the crystal-lithic tephras were performed at the U.S. Geological Survey (USGS) in Denver, Colorado. High-purity mineral separates together with standards were irradiated for 1 MW h in the central thimble position of the USGS TRIGA reactor using cadmium lining to prevent nucleogenic production of 40Ar. The neutron flux was monitored using Fish Canyon Tuff sanidine, using an age of 28.20 Ma ± 0.08 Ma (Kuiper et al., 2008) and isotopic production ratios were determined from irradiated CaF2 and KCl salts. For this irradiation, the following production values were measured: (36/37)Ca = 2.76 × 10−4 ± 0.29; (39/37)Ca = 7.12 × 10−4 ± 0.40; and (38/39)K = 1.29 × 10−2 ± 0.03. The irradiated samples and standards were loaded into 3 mm wells within a stainless-steel planchette attached to a fully automated ultrahigh-vacuum extraction line constructed of stainless steel. Samples were incrementally degassed and/or fused using a 20 W CO2 laser equipped with a beam homogenizing lens. The gas was expanded and purified by exposure to a cold finger maintained at –140 °C and two hot SAES GP50 getters. Following purification, the gas was expanded into a Mass Analyzer Products 215–50 mass spectrometer and argon isotopes were measured by peak jumping using an electron multiplier operated in analog mode. Data were acquired during 10 cycles and time zero intercepts were determined by best-fit regressions to the data. Ages were calculated from data that were corrected for mass discrimination, blanks, radioactive decay subsequent to irradiation, and interfering nucleogenic reactions.

The 40Ar/39Ar data are given in Table 1 and presented in graphical form in Figure 9. Two to three grain aliquots of amphibole from two horizons yielded 40Ar/39Ar plateau ages of 3.89 ± 0.17 (2σ) and 4.30 ± 0.10 (2σ) (Table 1). These ages are older than the presumed depositional age of the sediments, suggesting detrital input of older source material into the basin. Reworked material into the basin is also indicated by the 40Ar/39Ar analyses of K-feldspar and plagioclase. Thirteen grains of K-feldspars were individually fused from sample 09JTA-1143B and their 40Ar/39Ar ages range from 94.35 ± 0.52 (2σ) Ma to 3.33 ± 0.10 (2σ) Ma (Fig. 9A, Table 1). The K-feldspar is a mixture of microcline and sanidine. We interpret the ages greater than 4 Ma as reworked microcline, possibly from the Sierra Nevada batholith, and the 40Ar/39Ar ages ranging from 3.33 to 3.58 Ma as sanidine from local volcanic sources of the Basin and Range Province. Analytical uncertainties from individual K-feldspar 40Ar/39Ar ages are much less than the overall range in the ages, indicating that the sanidine was sourced from different eruptions over a period of ca. 250 ka. The youngest sanidine 40Ar/39Ar age of 3.33 ± 0.10 Ma from this sample places a maximum depositional age on the sample locality 09JTA-1143B.

Volcanic plagioclase grains from sample 09JTA-1142 yield 40Ar/39Ar ages ranging from 2.67 ± 0.42 (2σ) Ma to 3.88 ± 0.28 (2σ) Ma (Fig. 9B, Table 1). Like the K-feldspar data from sample 09JTA-1143B, the plagioclase from this horizon displays a range in 40Ar/39Ar ages that exceeds analytical uncertainties, consistent with multiple source inputs. The youngest plagioclase 40Ar/39Ar age of 2.67 ± 0.42 Ma places an upper limit on the age of sedimentation at this locality.

Mammal fossils document a still younger age for the uppermost sandy part of the Hunter Creek Sandstone (Fig. 3), in an interval above those beds dated isotopically. A collection of horse teeth from high in the section is dated as late Blancan (Blancan V) (Kelly and Secord, 2009). This collection is no older than 2.6 Ma, and possibly as young as 1.9 Ma or 1.5 Ma (Kelly, 2001; Kelly and Secord, 2009).

The interval of time from 5 to < 2.6 Ma saw the expansion, integration, and finally, filling of a larger stable lake system in the eastern part of the preserved and exposed Verdi basin. The lake lasted over 1 million years, based on bracketing tephra ages, and supported by accumulation rate and the thickness of the measured diatomite section. This lake preserved a thick accumulation of diatomite punctuated by thin tephras, mostly crystal-lithic in composition, from local volcanic sources. Lacustrine sediments can be traced laterally into deltaic facies in all directions where the sediments are preserved. These peripheral deltaic intervals are consistent with several small streams bringing sediment to the lake margins. The clean diatomite in the center of the small basin suggests that there were no large rivers bringing large influxes of mud and silt during seasonal flooding. In contrast, repeated clastic influxes did occur in the contemporary Honey Lake basin 80 km to the north, producing a distinctive section of turbidites and gravites (Trexler et al., 2009).

Sandy diatomite preserved in the northern and southern parts of the basin suggests that the lake was a closed basin not significantly bigger than the modern extent of Verdi basin lacustrine outcrops (maximum of ∼25 km2) (Fig. 2). Scattered small outcrops along the east front of the Carson Range to the southeast suggest the lake may also have occupied the western part of what is now the modern southwest Reno basin. Gravity modeling shows that the diatomite section does not extend down-dip into the subsurface much beyond the current extent of outcrop in west Reno (Widmer et al., 2007; Cashman et al., 2012). Deltaic sediments in the upper section indicate that the western edge of the lake was not far away and appears to have encroached into the lake with time (Fig. 4C). There is no evidence that the Sandstone of Hunter Creek was part of a larger lake linked to other Neogene basins in the region (e.g., the Boca basin to the west, Long Valley basin and Honey Lake basin to the north, or Gardnerville basin to the south) that might have contained lakes at the same time (e.g., Trexler et al., Mass et al., and Cashman et al. inOldow and Cashman, 2009).


Earliest Neogene Verdi basin sedimentation coincided with three tectonic events that directly affected the incipient basin between 11 and 10 Ma: (1) a shift from intermediate to mafic magmatic activity, followed by (2) cessation of local volcanism, and (3) generation of local fault-related topographic relief. The youngest local lavas and volcaniclastic syneruptive sediments (including lahars) are intercalated with the earliest basin sediments. This earliest interval in the basin history did not last long; coarse, high-energy deposits and interbedded volcanic rocks are no younger than the youngest basaltic andesite flows, ca. 10.1 Ma.

The stratigraphic evidence that the earliest basin sediments were associated with active faults and related structural relief is indirect, and is based on the nature and composition of the sediments. They are coarse rock-avalanche deposits, talus, and debrites, and locally contain abundant, large, granitic clasts. Although there is no direct evidence for the orientation of the faults that created the structural relief, complex mapped fault geometries and kinematic evidence of reactivation suggest that at least some of the postdepositional (i.e., Pleistocene and younger) faulting reactivated preexisting fault surfaces. In particular, the steeply dipping fault surfaces with generally north strikes tend to exhibit oblique-slip motion, which may indicate that these are controlled by the orientation of older surfaces that were reactivated in a new stress field. A similar interval of uplift and faulting is also suggested in the north-central Sierra, to the west, where normal offset on a NNW-striking fault system is bracketed between 13 and 8 Ma (e.g., Henry and Perkins, 2001).

Most of the Miocene–Pliocene Verdi basin records an interval of tectonic quiescence that lasted until the late Pliocene. Volcanism continued in the Sierran arc, as evidenced by coarse air-fall tephras and pumice- and scoria-pebble conglomerate. Ash beds from more distant sources, such as the Yellowstone hot spot, are also preserved in the Sandstone of Hunter Creek (Perkins et al., 1998). However, there are no lava flows or lahars above the lowest part of the section. Sedimentary provenance of sand was arc volcanic rocks, dominated by feldspar and lithic fragments with very little quartz (Queen, 2008). This finding corroborates the interpretation that crystalline rocks at the core of the Sierra Nevada were not a significant sediment source accessible to this basin. Depositional thickness based on gravity models suggests a basin depth of 1.2 km (Abbott and Louie, 2000) to less than a kilometer (Widmer et al., 2007), although depositional thickness measured down-dip approaches 2.6 km (Fig. 3). There are no internal unconformities, fault-related sediments, fanning of dips, or any other indications of tectonic activity during the interval from ca. 10 Ma to 2.6 Ma. We interpret this record as the passive and quiet burial of existing volcanic topography. As our companion paper indicates, this passive record ended quite dramatically less than 2.6 million years ago (Cashman et al., 2012).

In summary, the most important results from our study of the Verdi basin deposits are: (1) the new (young!) dates from these rocks, and (2) the absence of evidence for Sierran uplift or related tectonism. New 40Ar/39Ar analyses place an older age limit of 2.6 Ma on tephras near the top of a lacustrine section in the Verdi basin deposits. The lake, accumulating diatomite and occasional tephras, records stable conditions that persisted for over 1 Ma. Floral and faunal assemblages are not consistent with the presence of a topographic barrier to the west.

This research was funded in part by NSF grants EAR9815122 and EAR0001130 to Cashman and Trexler. Throughout this study, we have all benefitted from discussions with many students and colleagues at the University of Nevada, Reno (Department of Geological Sciences and Engineering, and Nevada Bureau of Mines and Geology) and elsewhere. In particular, we thank Larry Garside, Jim Faulds, and Chris Henry (Nevada Bureau of Mines and Geology), Chris Benedict (Washoe County Department of Water Resources), John Louie (University of Nevada Department of Geological Sciences and Engineering), and many graduate and undergraduate students who have looked at these rocks with us over the years. Tom Kelly was generous with his time and vertebrate paleontological skills. Mike Perkins shared his encyclopedic knowledge of tephrostratigraphy of the region. We thank Cathy Busby, Keith Putirka, Chris Holm-Denoma, and an anonymous reviewer for very helpful comments on earlier versions of the manuscript. Many alert colleagues and friends in Reno have found and directed us to new and often temporary exposures of Pliocene–Pleistocene rocks, which we appreciate. Trade names used in this study are for reference only and do not imply endorsement by the USGS.