Magmatism, migrating topography, and the transition from Sevier shortening to Basin and Range extension, western United States

The switch from Mesozoic shortening to Cenozoic extension in the western part of the North American Cordillera was a fundamental tectonic transition with implications for orogenic systems worldwide, but its causes remain hotly debated and poorly understood. The many models focusing on this time interval reveal disagreements regarding these basic questions: What were the pre-extensional crustal structure, crustal thickness, and resulting topography across the western United States? What was the causal mechanism for initiation of continental extension? What was the detailed timing of important tectonic events across this time span, and how do those events factor into our understanding of the causal mechanisms for the switch from shortening to extension?

The prevailing view argues that Mesozoic crustal thickening in the Sevier fold-and-thrust belt produced a high plateau (the “Nevadaplano”) across the region of the present-day Great Basin (Fig. 1; DeCelles, 2004). Although the concept of the Nevadaplano is widely accepted, there is little agreement regarding the timing and cause of plateau uplift (cf. Parsons et al., 1994; Mix et al., 2011; Cassel et al., 2018), its peak elevation (cf. Chase et al., 1998; Wolfe et al., 1998; Best et al., 2009; Cassel et al., 2012), whether or not “rugged topography” may have been present on the plateau (cf. Chamberlain et al., 2007; Henry et al., 2012; Bahadori et al., 2018), and the causes and timing of its inferred “collapse” (cf. Sonder et al., 1987; McQuarrie and Chase, 2000; Colgan and Henry, 2009; Wells et al., 2012; Lee et al., 2017).

Related to these broader questions, there are more detailed questions about the resulting topography across this region. Studies based on the pattern of Cenozoic ash-flow tuffs that filled paleovalleys (e.g., Best et al., 2013; Henry and John, 2013) indicate an approximately north-south–oriented divide that passed down the middle of central Nevada (Fig. 2), slightly west of the mostly older, Eocene Elko Basin (e.g., Haynes, 2003; Lund Snee et al., 2016; Camilleri et al., 2017) and the mostly Late Cretaceous and Eocene Sheep Pass Basin (e.g., Druschke et al., 2009a, 2009b). However, the divide has also been inferred to lie along the axis of the Mesozoic Sierra Nevada arc during the Late Cretaceous (Van Buer et al., 2009; Sharman et al., 2015). When, why, and how did this eastward shift of the divide occur? Was the shift related to the south-sweeping middle Cenozoic volcanism of the ignimbrite flareup in the retroarc region (Fig. 2; e.g., Armstrong and Ward, 1991; Christiansen and Yeats, 1992)? How did volcanism and the addition of large volumes of magma and heat affect retroarc topography?

Here, we integrated new and previously published geologic mapping and U-Pb detrital zircon geochronology from Late Cretaceous(?)–Neogene sedimentary and volcanic rocks of the Elko Basin of northeast Nevada, one of the only successions spanning large portions of the time between Cretaceous crustal thickening and Neogene Basin and Range extension (Figs. 1 and 2). We discuss the paleogeographic and tectonic significance of the stratigraphic succession and published stable isotope records in the Elko area in the context of others from throughout the region. We also compiled ages for the oldest volcanic or sedimentary material deposited in a network of east- and west-draining paleovalleys active across the Great Basin in middle Cenozoic time. Drawing upon all of this information, we present a revised view of the paleogeographic and tectonic evolution of the northern Great Basin from Late Cretaceous to Neogene time. Finally, we discuss the implications for estimates of crustal thickness and topography prior to Cenozoic extension, including impacts for the Nevadaplano model.

Late Cretaceous and Cenozoic strata in northeast Nevada, in the area of the Eocene Elko Basin (Figs. 1–3), unconformably overlie a thick (~10 km) Neoproterozoic to Triassic section deposited along the passive margin of western North America (e.g., Willden and Kistler, 1979; Colgan et al., 2010), which formed after Neoproterozoic and early Paleozoic rifting of the Rodinia supercontinent (e.g., Lund, 2008; Yonkee et al., 2014). West of the Elko area, there lie deep-marine rocks of the Roberts Mountains and Golconda allochthons (Fig. 1) that were thrust across the continental margin in the earliest Mississippian and the Permian–Triassic, respectively (e.g., Stewart, 1980). These relations and the location of the initial ⁸⁷Sr/⁸⁶Sr = 0.706 isopleth imply that western Nevada was underlain by oceanic crust and that the thick passive-margin succession at the study area in northeast Nevada was underlain by thinned continental crust (Tosdal et al., 2000). East-dipping subduction beneath the Cordilleran margin initiated as early as the Early Triassic (e.g., Saleeby et al., 2008). The study area, which was in the retroarc region of the Sierra Nevada arc, experienced two episodes of shortening and metamorphism during periods of increased arc magmatism, one in the Middle to Late Jurassic ca. 170–155 Ma and the second in the Late Cretaceous ca. 120–70 Ma (e.g., Dallmeyer et al., 1986; Miller and Gans, 1989; Thorman et al., 1991; Smith et al., 1993; McGrew and Snee, 1994; Thorman and Peterson, 2003; du Bray, 2007; Zuza et al., 2020). Cretaceous thrust faulting at the latitudes of northern and central Nevada was confined mostly to the Sevier fold-and-thrust belt to the east of the study area, with significantly less shortening represented by the Central Nevada thrust belt in central and southern Nevada (Fig. 1; Taylor et al., 2000; Di Fiori et al., 2020).

Following Late Cretaceous subduction-related arc magmatism, volcanism initiated again as part of the middle Cenozoic ignimbrite flareup, characterized by widespread caldera-forming volcanism that migrated southward across the Great Basin, passing through northeast Nevada (Fig. 2) ca. 42–36 Ma (Brooks et al., 1995; Ressel and Henry, 2006; Ryskamp et al., 2008; Henry and John, 2013; Lund Snee et al., 2016). Although extension initiated locally as early as Eocene time in certain areas (e.g., Henry et al., 2011; Miller et al., 2012; Wells et al., 2012), recognition is growing that the primary extension that affected topography and supracrustal rocks across the hinterland was an episode of rapid slip on high-angle normal faults that took place mostly during middle to late Miocene time and culminated in the present Basin and Range topography (Lund et al., 1993; Miller et al., 1999; Stockli et al., 2002; Henry, 2008; Colgan and Henry, 2009; Colgan et al., 2010; Colgan, 2013; Konstantinou and Miller, 2015; Lund Snee et al., 2016). This contribution presents new geochronologic and geological data obtained in the ancestral Elko Basin. We then examined the sedimentary record for the constraints it provides on the controversial and enigmatic history during the Late Cretaceous to the Miocene and, specifically, how the surface topography may have changed across this time span.

We mapped geologic units and tuffaceous beds (Fig. 4), measured stratigraphic sections (Fig. 5), and collected samples for U-Pb detrital zircon geochronology in Oligocene and Miocene successions in Huntington Valley and the eastern Carlin-Piñon Range, northeast Nevada (Fig. 3), in and near the ancestral Elko Basin (Figs. 1 and 2). Our work refines geologic mapping by Smith and Howard (1977), Smith and Ketner (1978), Lund Snee (2013), Lund Snee and Miller (2015), and Lund Snee et al. (2016), who provided detailed descriptions of those rocks. The GSA Supplemental Material¹ contains detailed analytical methods, and it presents the results of seven U-Pb detrital zircon analyses in addition to those presented by Lund Snee et al. (2016). Figure 5 shows the new and previous geochronology results within their stratigraphic framework.

Some of the same sections have previously been sampled for stable isotope analysis of calcite cements, limestone, and paleosols (Horton et al., 2004; Mix et al., 2011), and our work refines the age constraints and understanding of the depositional context for Neogene rocks within those sections. Figure 6 shows stable isotope data from prior studies placed in their revised temporal positions, with permissible depositional ages conservatively bounded by the full 2σ uncertainty ranges for the depositional age constraints (e.g., including the 2σ uncertainties for weighted mean ages). The additional supplemental material in EarthChem (https://doi.org/10.26022/IEDA/112062) includes tables containing lithologic details, sample localities, (maximum) depositional age information, and analytical data. Using these same methods, we used published geochronologic data to improve depositional age constraints for published stable isotope data that were obtained in older, Paleogene strata in the area (locations in Fig. 3), for which a number of contradictory depositional ages have been reported (cf. Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Mulch et al., 2015; Smith et al., 2017; Ibarra et al., 2021). In most cases, the permissible age bounds were not given in those studies, and sample ages were reported as being absolute when in fact they were maximum depositional ages (MDAs).

The Elko area is one of the only places in the Great Basin where well-preserved sedimentary and volcanic rocks spanning the Late Cretaceous(?) to Neogene are exposed over appreciable areas (e.g., Stewart, 1980). Because the term “Elko Basin,” sensu stricto, refers to the part of northeast Nevada where the Elko Formation was deposited in Eocene time (see Camilleri et al., 2017), we use the term “ancestral Elko Basin” to refer to this general area over a wider time span.

As shown in Figure 6, little deposition is documented between the Cretaceous and early Eocene in the ancestral Elko Basin region (Smith and Ketner, 1976; Fouch et al., 1979; Rahl et al., 2002; Haynes, 2003; Crafford, 2007; Henry et al., 2011; Lund Snee and Miller, 2015), reflecting the history of gradual erosion that prevailed throughout much of the hinterland from the peak of Late Cretaceous deformation and magmatism until the middle Eocene (e.g., Van Buer et al., 2009; Konstantinou et al., 2012). The first strata deposited during this time span within the Elko area were Late Cretaceous(?) to early Eocene(?) red beds and limestones (Figs. 3, 4C, 5, and 6), probably in isolated topographic lows, fault-bounded basins, or the bottoms of paleovalleys (Armstrong, 1968, 1972; Smith and Ketner, 1976; Gans and Miller, 1983; Van Buer et al., 2009; Konstantinou et al., 2012; Long, 2012; Lund Snee, 2013; Henry, 2018). These early deposits, as well as the overlying Elko Formation (Fig. 5), contain clast compositions and detrital zircon age distributions that reflect recycling from strata presently exposed beneath the Cenozoic unconformity (Druschke et al., 2011; Ruksznis, 2015; Lund Snee et al., 2016; Canada et al., 2020).

The more extensive middle to late Eocene Elko Formation (Figs. 5 and 6), which locally reaches thicknesses of ~850 m (Henry, 2008), consists of a broadly upward-fining succession of conglomerate, sandstone, siltstone, shale, clay, marl, and limestone (Smith and Ketner, 1976; Solomon et al., 1979; Moore et al., 1983; Server and Solomon, 1983; Ketner and Alpha, 1992; Haynes, 2003; Lund Snee and Miller, 2015; Smith et al., 2017). The presence of Cenozoic tuffaceous material and other volcanic detritus is a key factor that distinguishes the Elko Formation from older units (Smith and Ketner, 1976; Lund Snee et al., 2016). Deposition of the Elko Formation began ca. 46.1 Ma (Fig. 5), based on a U-Pb zircon age of an ash-fall tuff deposited near its base (Haynes, 2003). The end of Elko Formation deposition is tightly constrained at ca. 38.4 Ma, based on a U-Pb detrital zircon MDA of 37.9 ± 0.5 Ma from its upper stratigraphic levels within the eastern Carlin-Piñon Range, south of Robinson Mountain (Fig. 3; sample ELKO-2 of Lund Snee et al., 2016) and a minimum depositional age of 38.47 ± 0.15 Ma from ⁴⁰Ar/³⁹Ar plagioclase analysis on an overlying ash-flow tuff nearby (sample H10–45 of Henry et al., 2015) (see Fig. 5). These dates revise an estimate of 40.4 Ma for the end of Elko Formation deposition by Smith et al. (2017), and they are compatible with estimates of ca. 39–38 Ma by Haynes (2003) and Mulch et al. (2015). The initiation of shallow basin development recorded by the onset of Elko Formation deposition ca. 46–44 Ma, shortly before arrival of volcanism (see below), indicates that volcanism occurring to the north in southern Idaho may have been responsible for a change in topography and regional stress state. The development of accommodation for deposition at this time, following tens of millions of years with little or no sedimentation, suggests initiation of a mechanism such as normal faulting, development of sags or uplifts, and/or establishment of broad paleodrainages (e.g., Howard, 2003; Smith et al., 2017; Henry, 2018; this study). Lithofacies characterization and δ¹³C_carbonate, δ¹⁸O_carbonate, and δD_glass measurements suggest that the Elko Basin experienced a profound transition in depositional setting between deposition of the lower and upper Elko Formation ca. 40 Ma, as volcanism became more proximal (Fig. 2), although published interpretations are contradictory. Mulch et al. (2015) suggested that elevations increased and lake waters freshened within the Elko Basin around this time, whereas Smith et al. (2017) proposed an upward transition from freshwater fluvial-lacustrine to saline and anoxic profundal settings, with substantial uplift only after the end of Elko Formation deposition.

The ca. 38.5–36.8 Ma Robinson Mountain volcanic field (Ressel and Henry, 2006; Henry et al., 2015; Lund Snee et al., 2016) is a thick succession (>1 km) of rocks associated with the ignimbrite flareup on the eastern flanks of the Carlin-Piñon Range (Fig. 3), and it conformably overlies the Elko Formation (Smith and Ketner, 1976, 1978; Ressel and Henry, 2006; Ryskamp et al., 2008; Lund Snee and Miller, 2015). Following volcanism, sedimentation and faulting effectively ceased in this area (Fig. 6), with few exceptions, until Basin and Range faulting began in the middle Miocene (Henry et al., 2011; Lund Snee et al., 2016; this study). Any faulting that accompanied magmatism and crustal flow at depth might have been limited to the immediate areas above the metamorphic core complexes, where the surface record is incomplete due to younger faulting and erosion (Miller et al., 1999; Colgan et al., 2010; Konstantinou et al., 2013a; Lee et al., 2017). Lund Snee et al. (2016) documented 10°–15° tilting events following volcanism, between ca. 36.8 and 33.9 Ma and between ca. 31.1 and 24.4 Ma (better constrained by this study to between ca. 31.1 and 25.1 Ma as shown in Figs. 4B, 5, and 6), which they interpreted to represent local deformation associated with deeper crustal flow leading to surface adjustments by either faulting and/or doming adjacent to the developing Ruby Mountains–East Humboldt Range metamorphic core complex. Best and Christiansen (1991) made a similar interpretation for the limited and localized faulting that occurred during the ~10 m.y. following volcanism throughout the Great Basin.

An early phase of gradual extension within this generally quiescent interval may be represented by deposition of ca. 25 Ma and progressively younger fluvial-lacustrine material near the base of the Humboldt Formation (Figs. 3, 4A, and 6; Lund Snee et al., 2016; this study), as well as the possibly correlative late Oligocene and/or early Miocene sedimentary sequence of Clover Creek (McGrew and Snoke, 2015) near the East Humboldt Range (Fig. 1). Subsequently, rapid slip initiated on basin-bounding faults at 17–16 Ma (e.g., Colgan et al., 2010), represented in northeast Nevada by thick middle to late Miocene deposits of the Humboldt Formation. Near the Ruby Mountains–East Humboldt Range (Figs. 3, 4B–4D, and 5), the Humboldt Formation locally exceeds 4 km thickness (Satarugsa and Johnson, 2000), by far dwarfing the thicknesses of older Cenozoic units.

Here, we expand the above discussion to include rocks of Late Cretaceous to Neogene age deposited elsewhere in the Great Basin, primarily in the Copper, White Sage, and Sheep Pass Basins (Fig. 1). As can be seen in Figure 6, the histories of deposition and tectonism in these areas are broadly similar, characterized by limited and localized Late Cretaceous to Paleogene sedimentation and faulting (e.g., Gans and Miller, 1983; Best and Christiansen, 1991; Burchfiel et al., 1992; Van Buer et al., 2009; Henry et al., 2011; Konstantinou et al., 2012; Long, 2012; Henry and John, 2013) followed by volcanism, a subsequent hiatus in sedimentation, and then rapid sedimentation and tilting in the middle Miocene.

Regionally, the most significant deposition in the Late Cretaceous to Paleogene interval was in the greater Sheep Pass Basin of east-central Nevada (Fig. 1), where deposition of up to ~1200 m of Late Cretaceous–middle Eocene Sheep Pass Formation and late Eocene Stinking Spring Conglomerate (Fig. 6) was associated with potentially 4 km of normal slip on the northwest-dipping Ninemile fault system (Druschke et al., 2009a, 2009b). Sedimentation and tilting occurred elsewhere in east-central Nevada (Gans et al., 1989), including near the Snake Range metamorphic core complex (Figs. 1 and 6), but deposition was localized (Best and Christiansen, 1991).

As in the case of the Elko area described above, the limited pre-Miocene deposition that did take place elsewhere in the Great Basin occurred several millions of years before ignimbrite flareup volcanism (Fig. 6), suggesting that early magmatic processes could have prompted changes in topography. In the Copper Basin of northeast Nevada (Fig. 1), localized late Eocene–Oligocene deposition may have occurred due to basin development associated with normal fault slip shortly before and during nearby volcanism (Axelrod, 1966; Rahl et al., 2002). Alternatively, these units (and the nearby Elko Formation) may have been deposited into paleochannels, some of which could have been partially dammed by faulting (Henry, 2008, 2018). In the White Sage Basin of western Utah, ~150 m of early Eocene deposits experienced modest tilting in the middle Eocene before being blanketed ca. 40–39 Ma by ignimbrite flareup volcanic rocks (Potter et al., 1995; Dubiel et al., 1996). These data sets appear to preclude significant activity and offset along Late Cretaceous to Oligocene faults, consistent with prior compilations for the Great Basin (Best and Christiansen, 1991; Colgan and Henry, 2009; Henry et al., 2011; Henry and John, 2013).

In most places, sedimentation effectively ceased following volcanism, lasting at least through most of the Oligocene. Similar to the Ruby Mountains–East Humboldt Range metamorphic core complex, gradual sedimentation near the Snake Range metamorphic core complex reinitiated as early as the late Oligocene or earliest Miocene (Miller et al., 1999; Ruksznis, 2015). Nevertheless, by far the most significant episode of sedimentation, tilting, and uplift in the Great Basin occurred during middle Miocene time, ca. 17–16 Ma, resulting in region-wide development of deep (often >2 km) half-graben basins that filled rapidly with sediments (Noble, 1972; Stockli et al., 2002; Colgan, 2013).

Numerous studies have targeted strata in the ancestral Elko Basin for stable isotope analysis in efforts to understand regional paleoelevation and paleoclimate histories (Fig. 3). Stable isotope measurements from Elko area carbonates have served as key constraints in regional studies arguing that large and rapid negative shifts in δ¹⁸O, which proceeded southward across the western United States approximately synchronous with migrating middle Cenozoic volcanism, indicate simultaneous south-migrating topographic uplift (Horton et al., 2004; Davis et al., 2009b; Mix et al., 2011; Chamberlain et al., 2012), supporting a model initially proposed by Gans (1990). Specifically, these conclusions were based on a proposed ~7‰–10‰ decrease in δ¹⁸O values at ca. 50–47 Ma in southwestern Montana and eastern Idaho (Kent-Corson et al., 2006), followed by a decrease of up to ~15‰ in the Elko Basin, proposed to have occurred between ca. 40.2 and 39.4 Ma (Mulch et al., 2015), and finally an ~4‰ decrease after ca. 23 Ma in southern Nevada (Chamberlain et al., 2012). This timing is generally corroborated by comparable negative shifts of δ¹⁸O values ca. 44–40 Ma in foreland basin deposits east of the Sevier fold-and-thrust belt at the latitude of northeast Nevada (Fig. 1), which have been attributed to high-elevation catchment areas in the Sevier hinterland (e.g., Carroll et al., 2008; Davis et al., 2009a). In the ancestral Elko Basin, the decrease in δ¹⁸O values was interpreted to suggest 2.5 km of uplift occurring over <2 m.y. (Chamberlain et al., 2012). Mulch et al. (2015) subsequently suggested that a component of the up to 15‰ decrease in δ¹⁸O values in the Elko Basin should be attributed to climatic and diagenetic factors, including late Eocene global cooling.

More recently, however, the sedimentary age constraints, mean δ¹⁸O values, and interpretations of depositional environment underpinning some of these studies have changed, both for the ancestral Elko Basin (Lund Snee et al., 2016; Smith et al., 2017) and the Sage Creek Basin of southwest Montana (Kent-Corson et al., 2010; Schwartz et al., 2019). In southwest Montana, Kent-Corson et al. (2010) revised the magnitude of the negative shift in δ¹⁸O values from 7‰ to only 4‰–5‰ (where all δ¹⁸O values are relative to standard mean ocean water [SMOW]). The 4‰–5‰ shift was corroborated by Schwartz et al. (2019), who also established that it occurred rapidly at ca. 47 Ma, across a conformable stratigraphic boundary. The ~1 m.y. time interval for this shift in δ¹⁸O values may be too rapid to be explained by topographic changes associated with Challis and Absaroka volcanism, which reached southwest Montana ~5 m.y. earlier.

In the Elko area, Lund Snee et al. (2016) found that the rocks previously mapped as part of the Oligocene Indian Well Formation (Fig. 3) are mostly middle Miocene and younger in age and hence part of the Humboldt Formation. As a result, that study recommended abandoning the Indian Well Formation name. Because much of the marked decrease in δ¹⁸O values (an interpreted ~14‰–15‰ decrease from ~+29.1‰ to +14.4‰) occurred across this angular unconformity, the age revisions imply that the timing of the shift could have occurred anytime within a large window of time between 40 and 15.5 Ma. Our new geochronologic data and more conservative approach to constraining temporal bounds (Fig. 6) do not significantly narrow the time interval over which the δ¹⁸O shift occurred, but they do constrain the ages more rigorously and in greater detail than prior studies (e.g., Mix et al., 2011; Mulch et al., 2015; Lund Snee et al., 2016; Smith et al., 2017; Ibarra et al., 2021). The record of δ¹⁸O values with improved age constraints shown in Figure 6 indicates considerable scatter in δ¹⁸O values both in the upper part of the Eocene Elko Formation and in lower levels of the Humboldt Formation for which stable isotope values have been measured. This scatter leads us to interpret a slightly different shift of ~–12‰, from ~+25‰ (but ranging between ~+14‰ and +30‰) in upper Elko Formation strata with preferred ages spanning ca. 40.9–38.6 Ma to ~+13‰ (ranging between ~+9‰ and +20‰) in strata within the lower Humboldt Formation with preferred ages spanning ca. 15.8–15.5 Ma. Although prior workers (e.g., Mulch et al., 2015) argued that the decrease in δ¹⁸O values occurred within the upper Elko Formation because of an ~14‰–15‰ decrease that is observed within that succession, we point out that this interpretation was made on the basis of three nonsequential data points and that δ¹⁸O values increase again by ~10‰ (from +14.4‰–+17.6‰ to +24.9‰–+26.1‰) immediately above and still within the Elko Formation (Fig. 6). Such rapid oscillation of δ¹⁸O values within a narrow part of the succession is unlikely to reflect changes in topography. Hence, we conclude that the ~12‰ (or possibly less) decrease in δ¹⁸O values observed in the Elko Basin occurred at an unknown rate sometime between ca. 38.6 and 15.8 Ma. Moreover, our new age constraints and stratigraphic thickness measurements show that >400 m of stratigraphic section are present below the lowest measured Humboldt Formation δ¹⁸O values, with a depositional age of 25.1 ± 0.2 Ma now established for a tuffaceous bed near the base of that unit (Figs. 4 and 5). A consequence of the improved age constraints is that we do not know when the decrease in Elko Basin δ¹⁸O values occurred relative to the onset of volcanism (Figs. 2 and 3) ca. 39–38 Ma (Ressel and Henry, 2006; Henry et al., 2015; Lund Snee et al., 2016). The sign and magnitude of the decrease in δ¹⁸O values are clearly consistent with an elevation increase, although not definitive, because of the potential for climatic and diagenetic influences over this interval.

Stable isotopic studies focusing instead on volcanic rocks that erupted across the Great Basin (including the Elko region) and westward to the Sierra Nevada flank have argued that a high plateau persisted across the former Sevier hinterland between 41 and 23 Ma; this argument is based on relatively low δD values in altered volcanic glass samples that span this age spectrum (Cassel et al., 2009, 2014, 2018). These results were interpreted to suggest that the inferred high plateau developed in Cretaceous time (Henry et al., 2012). However, the δD values vary with sample age and location. This led Cassel et al. (2018) to propose that elevations across the hinterland were ~2.25–3.0 km during late Eocene time and then fell by ~0.5–1 km by early Oligocene time, followed by ~1.5 km of uplift between the early and late Oligocene (to as high as 3.5 km in central Nevada), before eventually falling to present-day mean elevations around 1.75 km. Although the presence of an elevated plateau following the arrival of volcanism is consistent with the other geologic and stable isotopic evidence discussed here, these proposed oscillatory elevation changes are difficult to reconcile with geologic evidence as they imply multiple episodes of faulting and sedimentation that are not corroborated by the geologic record (see above and summaries by Best and Christiansen, 1991; Henry et al., 2011; Henry and John, 2013). Moreover, the age brackets of δ¹⁸O values from Elko Basin carbonates (Fig. 6; Horton et al., 2004; Mix et al., 2011; Chamberlain et al., 2012; Mulch et al., 2015) are not consistent with models that require elevations to have decreased between ca. 42 and 13 Ma (e.g., Coney and Harms, 1984; Sonder et al., 1987; Bahadori et al., 2018; Cassel et al., 2018), because δ¹⁸O and δ¹³C values were generally low from ca. 16 Ma onward, and prior to ca. 42 Ma, they were substantially higher (Fig. 6).

In summary, recent work to improve age constraints and better define depositional settings in both northeast Nevada and southwest Montana (Kent-Corson et al., 2010; Lund Snee et al., 2016; Schwartz et al., 2019; this study) indicates that the stable isotope record does not alone provide clear evidence for south-migrating topography across the region. On the basis of more rigorous age constraints, the negative shift in δ¹⁸O values within the ancestral Elko Basin cannot be decisively linked to the age of south-migrating volcanism, although it is consistent with uplift occurring at that time or later and is inconsistent with models requiring a decrease in elevation anytime between ca. 42 and 13 Ma (Fig. 6). Far stronger evidence for south-migrating topography is demonstrated by the other data sets discussed here: the record of sedimentation, faulting, erosion, and reorganization of drainage networks. Stable isotope studies based on δD values in altered volcanic glass present a more complex pattern with space and time. Published interpretations of those results cannot presently be reconciled with the existing record of sedimentation and faulting.

Rocks preserved within a mapped network of Eocene–Oligocene paleovalleys (Fig. 2) provide additional insights for the region’s topographic evolution that can be interpreted in concert with the sedimentary record. The paleovalleys are defined and mapped by their thicker sequences of volcanic fill and can be traced generally east-west from range to range across Nevada (Henry, 2008). The rocks filling the paleovalleys are precisely dated and consist of well-correlated middle Cenozoic ignimbrites that flowed hundreds of kilometers east and west along these channels from their source calderas (e.g., Henry and John, 2013). The flow directions defined by these relations outline a north-south–trending paleodivide through central Nevada (Henry, 2008; Best et al., 2013), which has been widely displayed in subsequent publications portraying the paleotopography of the west. The paleodivide is viewed as a static feature at least through the early Cenozoic, if not as far back as the Cretaceous (e.g., Henry et al., 2012).

Figure 2 shows the Cenozoic temporal evolution of volcanism together with the ages of the oldest material dated within each paleodrainage (compiled from MacGinitie, 1941; Yeend, 1974; Goldstrand, 1992, 1994; Garside et al., 2005; Henry, 2008; Henry et al., 2012; Henry and John, 2013; Dumitru et al., 2015, 2016). The age of the oldest dated material in each paleovalley decreases systematically southward, broadly accompanied by the southward progression of volcanism. We consider it unusual that no basal paleovalley deposits have known ages significantly older than nearby magmatism. This is especially remarkable considering that the ignimbrites were capable of traveling hundreds of kilometers north and south of their eruptive centers (Henry and John, 2013). For each stage of migrating volcanism, the earliest deposits are preserved near the source calderas, filling paleovalleys to the west, east, and sometimes north. Yet—critically—the record of fill preserved at the bottoms of paleovalleys (Fig. 2) shows few to no examples of eruptive products that would have had to travel significantly south (e.g., ≥100 km).

The likeliest explanation is that most or all of these drainages did not exist more than a few million years before magmatism began at a given latitude and paleovalleys were filled with volcanic material. We propose that the paleovalleys could have developed diachronously in response to dynamic topographic uplift that likely occurred during south-migrating ignimbrite flareup volcanism, when large amounts of magmatic material and thermal energy were added to the crust. In this context, fluvial paleovalleys with valley-margin relief approaching 1.2 km (e.g., Henry et al., 2012) were incised into the rising hinterland, similar to the paleovalleys that formed during the Paleocene–Eocene to the north of the study area (e.g., Schwartz and Schwartz, 2013; Schwartz et al., 2019). Migrating eruptions would have progressively filled any newly developed valleys with resistant volcanic rocks that eventually (because they blanketed the landscape) constructed an elevated, relatively flat plateau such as that described by Best et al. (2009), with surface elevations substantially increased by magmatic and thermal input into the crust.

Figure 7 presents a schematic diagram of the proposed paleotopographic evolution ca. 36 Ma, immediately following the end of volcanism in the Elko area. As magmatism swept south across the Sevier hinterland in Eocene and Oligocene time, it prompted the growth of topography due to the voluminous addition of magma to the crust together with accompanying thermal uplift. We suggest that highlands developed near the main eruptive centers at any given time and that the approximately north-south–trending paleodivide proposed by Henry (2008), Henry et al. (2012), and Best et al. (2013) was a southward-propagating dynamic feature, with the highest topography located above the region of the most active caldera centers. Prior to volcanism, drainages likely flowed east and west away from the axis of the Cretaceous arc (Figs. 2 and 6), as defined by the locus of intrusion of the youngest plutonic complexes (Van Buer et al., 2009; Van Buer and Miller, 2010; Sharman et al., 2015). This inference is based on the significantly higher calculated magnitudes of erosion (~5–7 km) in the arc compared to values of 1–3 km across the back-arc region west of the Sevier fold-and-thrust belt (Van Buer et al., 2009). The detrital zircon signatures of Late Cretaceous–Eocene sediments deposited in the California forearc basin, west of the arc axis, also indicate derivation from the Sierra Nevada magmatic arc (Figs. 1 and 2) and do not indicate detectable sediment derivation from farther east (Sharman et al., 2015). Some drainages may also have flowed southward toward the Mojave region (Fig. 1), much of which was a marine environment during Paleocene time (e.g., Lofgren et al., 2008; Lechler and Niemi, 2011). Evolving southward-moving uplift would have tended to reorganize preexisting topography and drainages that were holdovers from the Late Cretaceous, resulting in paleodrainages that emanated to the west, east, and south from the new paleodivide in Nevada (e.g., Best et al., 2013; Henry and John, 2013; Lechler and Niemi, 2011; Miller et al., this volume). Notably, exposures of the Eocene Titus Canyon Formation in the Funeral Range, southeast California (Fig. 1), herald the onset of volcanism in north-central Nevada and include clast compositions and detrital zircon age components that suggest that sediment traveled through this region in south-flowing river systems from higher elevations located at least 300 km to the north-northeast (Miller et al., 2019; Miller et al., this volume). At the same time, paleodrainages were less likely to flow northward from active calderas due to the presence of predecessor high topography (Fig. 7).

Removal of the Farallon slab during middle Cenozoic time is thought to have caused the asthenospheric upwelling and subsequent ignimbrite flareup magmatism that would have also thinned the lithosphere below the Great Basin (Humphreys, 1995). Thermally driven uplift is an established consequence of both lithospheric thinning and addition of magmatic material (e.g., Lachenbruch and Morgan, 1990). The precise magnitude of elevation increase due to ignimbrite flareup volcanism is uncertain, but analogy with other regions suggests that it was 1 km or more (see Crough, 1978; Pierce et al., 1992; Larimer et al., 2019; Schwartz et al., 2019). Along and adjacent to the Yellowstone hotspot track (approximately the Snake River Plain on Fig. 1), localized volcanism between middle Miocene time and the present was accompanied by an east-migrating zone of pronounced uplift, faulting, reorganization of drainage systems, and shifts of the continental divide (Anders and Sleep, 1992; Pierce et al., 1992; Beranek et al., 2006; Coble and Mahood, 2016; Camilleri et al., 2017; Larimer et al., 2019). Detrital zircon records indicate that northeastward migration of Yellowstone hotspot magmatism profoundly disrupted drainage networks, causing streams to emanate away from the associated topographic bulge as it progressed, producing a northeast-oriented paleodivide that was centered either on or at the southern margin of the hotspot track, which later evolved into a crescent-shaped divide around the northwest, east, and southeast sides of Yellowstone (Beranek et al., 2006; Camilleri et al., 2017). The uplift near Yellowstone has led to 1 km deep river incision since 3.6 Ma, an incision rate of nearly 300 m/m.y. (Pierce and Morgan, 2009). If the middle Cenozoic Great Basin paleovalleys shown in Figure 2 developed as a result of the magmatism as we propose, then the up to 1.2 km of mapped paleovalley relief (Henry, 2008; Henry et al., 2012) provides a minimum estimate for the amount of synvolcanic uplift, comparable to recent incision near Yellowstone. The lack of evidence for paleorivers that incised into this elevated volcanic tableland following the decline of volcanism (Figs. 6 and 7) might have been a consequence of postvolcanic areas being slightly lower than areas to the south that were still experiencing active volcanism, and/or this might have been due to the erosional resistivity of the volcanic rocks. The uplift rates along the Yellowstone hotspot track not only demonstrate the significance of thermally driven topographic changes but also underscore that hundreds of meters of paleovalley erosion/incision may occur relatively rapidly, such as on the order of 1–10 m.y., rather than over long periods of geologic time, such as from the Late Cretaceous to middle Eocene (see, e.g., Colgan and Henry, 2017). We propose that the same processes that are active today near Yellowstone were active in the Great Basin, but at a grander scale, during the much more voluminous middle Cenozoic ignimbrite flareup.

Based on the evidence presented above from the sedimentary, structural, and magmatic records of the Elko area, integrated with data from surrounding areas, we present a summary view of the evolution of the Late Cretaceous to Cenozoic paleogeography of the northern Great Basin, as illustrated in Figure 8. The time line applies to northeast Nevada (right-hand panels of Fig. 6), but we suggest that it is also applicable to much of the Great Basin, especially pertaining to regional tectonic events (left-hand panels of Fig. 6). This paleogeographic and tectonic history is based upon the well-preserved and less controversial surface geologic record, but it is also intended to reconcile some of the contradictory models for the tectonic evolution of the region.

Surface-breaking thrust faults were active along the Sevier belt (Fig. 1) in the Late Cretaceous (Fig. 8A), and they created topography as the thrust belt finalized its development and shed its erosional debris into the foreland basin (e.g., Malone et al., this volume). In addition to voluminous foreland basin deposits, erosion of the Sevier belt is reflected by the basal Cenozoic unconformity map, which shows erosion carved much deeper into the miogeoclinal section within the thrust belt (e.g., Armstrong, 1968, 1972; Van Buer et al., 2009; Konstantinou et al., 2012). The Sevier thrust faults ceased most of their motion by Paleocene time (Fig. 6) as deformation moved east to the Rocky Mountains during the middle Late Cretaceous to Eocene Laramide orogeny (e.g., DeCelles and Coogan, 2006; Copeland et al., 2017), but motion along the Paris thrust of the Sevier belt in southeast Idaho and northeast Utah may have continued into the Oligocene (Malone et al., this volume). The switch to Laramide-style deformation is thought to have been linked to the onset of shallow slab subduction of part of the Farallon plate (Fig. 8A), as documented by both the cessation of magmatism in the Sierra Nevada arc and the eastward shift of deformation (e.g., Dickinson and Snyder, 1978).

Throughout the Late Cretaceous and early Cenozoic (Fig. 6), northeast Nevada and much of the Great Basin region to the west of the Sevier thrust belt (its hinterland) appear to have experienced only modest erosion (<3 km in most areas), with evidence for faulting and tectonism in the surficial record only located around a handful of previously identified structures (e.g., Van Buer et al., 2009; Konstantinou et al., 2012; Long, 2012). These inferences are further supported by reconstructions of the middle Cenozoic unconformity and paleogeologic maps of units beneath that unconformity. The observation of generally low conodont alteration index values across much of the Great Basin (Harris et al., 1980; Gans and Miller, 1983; Gans et al., 1990; Crafford and Harris, 2005) likewise indicates burial only to stratigraphic depths and is consistent with erosion mostly limited to the upper part of the Paleozoic–Mesozoic shelf stratigraphic section. During the Late Cretaceous and early Cenozoic, multiple lines of evidence discussed above suggest that the regional topographic divide was located near the axis of the Cretaceous magmatic arc (Figs. 2, 6, and 8A; Van Buer et al., 2009; Sharman et al., 2015). To the east of the Sevier fold-and-thrust belt, parts of the foreland basin system in Utah were near sea level (Fig. 8A) in the Late Cretaceous (DeCelles and Coogan, 2006). The southern margin of the Great Basin also lay near sea level in Paleocene and possibly early Eocene time (Figs. 1 and 6), as indicated by Paleogene marine fossils near the southeastern part of the Sierra Nevada, on the northern margin of the Mojave Desert (Lofgren et al., 2008; Lechler and Niemi, 2011). Consequently, elevations most likely decreased both to the east (Fig. 8A) and south from the Cretaceous Sierra Nevada arc and the Sevier hinterland.

The Elko and Copper Basins developed during middle Eocene time as shallow basins in northeast Nevada (Figs. 1 and 2), initiating by ca. 46 Ma and perhaps locally as early as ca. 49 Ma (Haynes, 2003; Lund Snee et al., 2016; Smith et al., 2017). At about this time, the Farallon shallow slab is inferred to have started to steepen (Fig. 8B), triggering upwelling of hot asthenosphere that contributed to an influx of magma and heat to the crust (e.g., Armstrong and Ward, 1991; Humphreys, 1995; Konstantinou et al., 2012; Konstantinou and Miller, 2015). The Elko and Copper Basins (Fig. 1) provide the first indications of an eastward shift of the topographic divide (Fig. 8B) in middle Eocene time. Deposits preserved in paleovalleys indicate both eastward and westward flow of ignimbrites away from an area west and north of the Elko Basin (Fig. 2) in northern Nevada (Henry, 2008). As suggested above, the mapped paleodrainages developed approximately during magmatism (Fig. 6), changing an earlier landscape through topographic growth and reorganization and replacement of preexisting drainage networks (Fig. 7).

Basin development may have been a result of normal faulting (Vandervoort and Schmitt, 1990; Rahl et al., 2002; Haynes, 2003; Howard, 2003) and/or isostatic adjustments associated with steepening of the Farallon slab and asthenospheric upwelling (Smith et al., 2017) and/or the onset of magma chamber formation and volcanism. Faulting before and during ignimbrite flareup volcanism “was minor and/or exceedingly local” throughout the Great Basin (e.g., Henry and John, 2013, p. 954), but such faulting shortly before—and in rare cases during—volcanism is observed in several areas, including northeast Nevada (e.g., Henry et al., 2011) and central Nevada (e.g., Gans et al., 1989; Best and Christiansen, 1991; Miller et al., 1999; Druschke et al., 2009a; Ruksznis, 2015). Localized faulting may have occurred due to rapid but differential emplacement of voluminous magma bodies and heat transfer to the upper crust and/or localized thermal weakening of crust experiencing far-field extensional strain, as has been suggested for the vicinity of all three Great Basin metamorphic core complexes (Miller et al., 1999; Konstantinou et al., 2012, 2013a; Konstantinou and Miller, 2015; Lund Snee et al., 2016; Lee et al., 2017). Arrival of volcanism was accompanied by a marked increase of heat input to the crust, as indicated by increased high-temperature mineral (zircon and monazite) growth ca. 42 Ma at deeper levels of the crust now exposed in the Ruby Mountains–East Humboldt Range metamorphic core complex (Fig. 6; Howard et al., 2011).

Sedimentation effectively ceased with the inception of volcanism (see above and Fig. 6), which blanketed the focus area with volcanic flows and ignimbrites beginning ca. 38 Ma (Haynes, 2003; Ressel and Henry, 2006; Lund Snee et al., 2016). These volcanic rocks formed a flat regional tableland (Best et al., 2009) that persisted in northeast Nevada with little erosion until the onset of Neogene Basin and Range faulting. Similar relationships are likely across the entire northern Great Basin (Fig. 6), as documented by the widespread preservation of volcanic rocks above the regional lower Cenozoic unconformity (e.g., Gans and Miller, 1983; Van Buer et al., 2009; Konstantinou et al., 2012). Following the last eruption within the study area ca. 36.8 Ma (Lund Snee et al., 2016), magmatism continued southward (e.g., Ryskamp et al., 2008), likely in tandem with elevation gain, as proposed in this paper (Fig. 7).

Partial melting and magmatism continued in the deeper crust within the developing Ruby Mountains–East Humboldt Range metamorphic core complex and the Albion–Raft River–Grouse Creek metamorphic core complex to the north (Figs. 1, 2, and 6), long after the cessation of surface volcanism, and concomitant with relative rise of metamorphic rocks (McGrew and Snee, 1994; McGrew et al., 2000; Howard et al., 2011; Konstantinou et al., 2013a). The persistence of elevated temperatures, combined with ongoing and prior magmatic addition to the crust from the mantle, likely ensured that topography remained thermally elevated to some degree (Fig. 7), at least through much of the Oligocene. The near absence of sedimentary deposits between ca. 38 and 25 Ma in the study area (Fig. 6), and following volcanism throughout the northern Great Basin in general (Henry et al., 2011), confirms that whatever surface-breaking faulting that occurred during the 10 m.y. or more time span following volcanism was very limited in extent and magnitude.

The sedimentary record in northeast Nevada indicates that tectonic quiescence and gradual erosion occurred between latest Oligocene and middle Miocene time (Fig. 6). Locally, however, lacustrine sedimentation (indicating the formation of basin accommodation) may have initiated during this time (Figs. 4B, 5, and 6), based on the ages of the earliest fluvial-lacustrine sediments deposited above Eocene and Oligocene volcanic rocks in Huntington Valley (Figs. 4B and 5) and near the East Humboldt Range (Fig. 1; Frerichs and Pekarek, 1994; McGrew and Snoke, 2015; Lund Snee et al., 2016). Minor deposition in the latest Oligocene or early Miocene has also been recorded near the Snake Range metamorphic core complex (Gans et al., 1989; Miller et al., 1999; Ruksznis, 2015) and elsewhere (Fig. 6).

As discussed above, rapid slip on Basin and Range normal faults, with formation of ensuing topography similar to that of today, occurred across most of the central part of the northern Basin and Range Province (Figs. 1, 2, 6, and 8C) ca. 17–16 Ma (e.g., Noble, 1972; Lund et al., 1993; Miller et al., 1999; Stockli, 2005; Colgan et al., 2010). Across this region, fault slip rates decreased beginning ca. 12–10 Ma (Fig. 6), based on thermochronologic data and sedimentation rates (e.g., Colgan et al., 2008; Colgan and Henry, 2009). Extension that began in these central areas subsequently propagated west, east, and north (Surpless et al., 2002; Stockli, 2005; Colgan et al., 2006; Lerch et al., 2008). Extension continues at a slow rate today as active slip takes place primarily on faults now close to the boundaries of the province (Thatcher et al., 1999; Kreemer et al., 2010). The timing of rapid extension coincided closely with a number of notable tectonic events (Figs. 6 and 8C), including final removal of the Farallon slab ca. 20 Ma (Humphreys, 1995), development of a gap in the Farallon slab ca. 17 Ma and subsequent impingement of the Yellowstone hotspot (Liu and Stegman, 2012), and the progressive development of the San Andreas transform boundary with northward migration of the Mendocino triple junction over Neogene time (Atwater and Stock, 1998).

The paleogeographic and crustal history outlined in this synthesis (Figs. 6, 7, and 8) has direct implications for the topographic and crustal evolution of the hinterland region between the Cretaceous arc and Sevier thrust belt (Figs. 1 and 2). Multiple lines of evidence suggest that appreciable elevation gain (probably 1 km or more, as suggested above) may have taken place much later than Mesozoic time (Fig. 8A), roughly synchronous with and persisting to some degree after Cenozoic volcanism swept across the region (Figs. 2 and 8B), before subsiding to the present ~1.5–2.0 km average elevations and sawtooth topography during and/or after Basin and Range extension (Fig. 8C). The growing database discussed here indicates an important time lag between crustal thickening and extension that is inconsistent with suggestions that a high plateau was supported by gravitationally unstable crust overthickened during the Mesozoic (e.g., Sonder et al., 1987; Chase et al., 1998; Druschke et al., 2009b; Wells et al., 2012; Wells and Hoisch, 2012; Craddock Affinati et al., this volume). In addition, an important consideration is that the heat budget represented by ultimately mantle-derived Cenozoic magmatism by far exceeded that related to thermal equilibration of crust thickened by thrust faulting (Gottlieb, 2018; Gottlieb et al., this volume).

The thermal effects from voluminous and widespread ignimbrite flareup magmatism are rarely considered in topographic and tectonic models of the Great Basin. Studies that estimate preextensional crustal thicknesses by restoring Cenozoic extension (e.g., Bahadori et al., 2018; Long, 2018) do not account for the substantial thicknesses of middle Cenozoic volcanic material added to the surface in many areas nor the potentially much greater volumes of associated plutonic material; thus, their inferred postthickening and preextensional crustal thicknesses are likely overestimates. A (map view) restoration of topography across the northern Great Basin by Bahadori et al. (2018) proposed a narrow, tall (crest ≥4 km and peaks >6 km) mountain range atop an ~55–60-km-thick welt of crust in the Eocene. That study restored preextensional crustal thicknesses based on the kinematic model by McQuarrie and Wernicke (2005), combined with an isostatic compensation model. The geometry of the resulting crustal welt is broadly similar to that shown by Long (2018), which was based largely on the Sevier thrust belt reconstruction of DeCelles and Coogan (2006). The mountain chain modeled by Bahadori et al. (2018) lies ~200 km east of the middle Cenozoic paleodivide that was inferred by Henry et al. (2012) and Best et al. (2013) using the paleoflow directions of channelized ignimbrites. The suggestion of a rugged, ≥4-km-tall mountain chain along the Utah-Nevada border, supported by relatively thick crust, is at odds with evidence that prevolcanic erosion magnitudes were modest and smoothly distributed throughout the area of the inferred crustal welt (Gans and Miller, 1983; Miller and Gans, 1989; Konstantinou et al., 2012; Long, 2012). It would also be unusual for such steep topography to develop just to the west of the active fold-and-thrust belt. This disagreement between these estimates of crustal thickness and topography and the geologic data described in this paper suggests the need for (1) retrodeformation studies that consider the limited crustal thicknesses that lay beneath the passive-margin sequence west of its depositional hinge line; (2) tighter constraints on the magnitude of westward crustal underthrusting (see Craddock et al., this volume; Gottlieb et al., this volume); (3) incorporation of updated models of Cenozoic extension and possible magmatic additions to the crust during the ignimbrite flareup; and (4) consideration of the thermal state of the crust and the likelihood of regional-scale lower-crustal flow that might have flattened the Moho before and during extension (Gans, 1987).

The assortment of geologic data presented here is incompatible with suggestions that crustal thicknesses became so great during Mesozoic shortening that they led to gravitationally driven extensional collapse (e.g., Wells et al., 2012). There is also little evidence in the record of sedimentation, stable isotope values, and deformation, such as surface faulting, to suggest significant changes in elevation between the Late Cretaceous and the time shortly before the arrival of middle Cenozoic volcanism. What little deposition occurred was mostly within the Sheep Pass Basin, where up to ~1 km of Sheep Pass Formation sediments was deposited throughout Late Cretaceous to middle Eocene time (Figs. 1, 2, and 6; Druschke et al., 2009a, 2009b). We point out that the gradual, localized occurrence of normal faulting thought to have provided accommodation for these deposits does not necessarily signify wholesale gravitational collapse of overthickened crust across the immense region envisaged as encompassing the Nevadaplano. Moreover, topographic relief (which would result from widespread surface-breaking faulting) likely was low across most of the Sevier hinterland before middle Cenozoic time (Fig. 8A), based on the depositional patterns of far-traveled Cenozoic ash-flow tuffs (e.g., Best et al., 2009) and the modest magnitudes of pre-Eocene erosion and tilting documented in the Elko region (Brooks et al., 1995; Henry et al., 2011; Lund Snee et al., 2016; Canada et al., 2020; this study) and across the hinterland in general (Gans and Miller, 1983; Gans et al., 1990; Crafford and Harris, 2005; Van Buer et al., 2009; Long, 2012; Konstantinou et al., 2013b). Suggestions that the hinterland was early on characterized by rugged, mountainous topography (Druschke et al., 2011; Bahadori et al., 2018; Bahadori and Holt, 2019) are clearly at odds with the above set of observations.

Few constraints are available for absolute elevations of the Sevier hinterland prior to extension, during the Late Cretaceous and early Cenozoic. Measurements from Eocene fossil leaves in Copper Basin (Figs. 1 and 2), representing the time at the onset of volcanism, provide widely distributed elevation estimates ranging from 0.6 to 1.2 km (Christiansen and Yeats, 1992) and 1.6 ± 1.6 km (Chase et al., 1998) to 2.0 ± 0.2 km (Wolfe et al., 1998) and 2.8 ± 1.8 km (Chase et al., 1998). This broad range complicates efforts to employ such estimates quantitatively. A definitive minimum hinterland elevation bound of 1.2 km was provided by Henry et al. (2012), based on measured middle Cenozoic paleovalley depths, which we suggest developed only after uplift related to ignimbrite flareup volcanism. Probably the most reliable estimates of absolute elevation are provided by two clumped isotope studies. Snell et al. (2014) estimated that absolute elevations in the Sheep Pass Basin (Fig. 1) of east-central Nevada ranged between 2.0 and 3.1 km in latest Cretaceous to early Paleocene time. Also using clumped isotope thermometry, Lechler et al. (2013) estimated only ≤2 km paleoelevation for the Sheep Pass Basin, integrated over the younger but overlapping latest Cretaceous–early Eocene interval. Although different, these two elevation estimates overlap at ~2 km, suggesting that this may be a reasonable elevation value for east-central Nevada in latest Cretaceous–early Eocene time. Additional lines of evidence support suggestions of only modest elevations across the Sevier hinterland prior to volcanism. As noted, marine fossils, stable isotope data, and detrital populations show that the southern Sierra Nevada and areas slightly to the east (~35.5°N latitude), which have a similar geologic history to the Great Basin, were at or near sea level in the Paleocene (Figs. 1 and 6) and may have remained very low (<1 km) into Eocene time (Lofgren et al., 2008; Lechler and Niemi, 2011). If a significantly elevated plateau was present across the Great Basin in Late Cretaceous and Paleogene time, then it must have been limited to areas north of ~37°N–38°N and bounded on the south by slopes leading nearly to sea level.

It is challenging to reconcile the above narrative of Great Basin surface evolution with the implications of 10–20 km of relative uplift implied by quantitative geobarometry on Cretaceous metamorphic assemblages in metamorphic core complexes (Hodges et al., 1992; Lewis et al., 1999; McGrew et al., 2000; Cooper et al., 2010; Hallett and Spear, 2013). These analyses represent the primary evidence supporting suggestions of large crustal thicknesses that drove gravitational collapse of the hinterland prior to Miocene time. A more complete discussion of these questions is provided by Hoiland et al. (this volume), but two hypotheses are relevant: (1) In recent years, it was recognized that the assumptions underpinning these geobarometric methods may in some cases be invalid, particularly the expectations that mineral assemblages were in equilibrium during formation (Spear et al., 2014) and that the pressure measurements can be interpreted as representing steady-state overburden pressures (proportional to burial depth) rather than transient and/or nonisostatic (tectonic) stresses (Schmalholz et al., 2014; Gerya, 2015). This leaves open a number of other possibilities to explain high pressure estimates in Great Basin metamorphic core complexes, including tectonic “overpressure” (Henry et al., 2018; Thorman et al., 2020; Zuza et al., 2020; Hoiland et al., this volume). (2) If some or all of the proposed uplift in developing metamorphic core complexes was Cenozoic in age, then uplift could have occurred with little surface-breaking extension provided that lower-crustal rocks were locally decoupled from surface deformation due to strongly elevated heat flow during midcrustal melting (MacCready et al., 1997; Miller et al., 1999; Konstantinou et al., 2012; Lund Snee et al., 2016; Lee et al., 2017). This mechanism would explain the difference in timing of subsurface uplift versus surface-breaking extension, as elegantly demonstrated for the Albion–Raft River–Grouse Creek (Fig. 1) metamorphic core complex (Konstantinou et al., 2013a).

We have presented an updated view of the enigmatic transition from Mesozoic shortening to Cenozoic extension in the Great Basin (Figs. 1 and 2), focusing primarily on the supracrustal records of sedimentation, erosion, faulting, volcanism, and stable isotope values, and the ways in which they relate to topography development. This integrated record shows that gradual erosion, limited deposition, and general tectonic quiescence prevailed between Late Cretaceous and middle Cenozoic time (Fig. 6). Although surface-breaking faults are documented across this time interval, they were local in scale and significance, involving relatively low magnitudes of slip (e.g., Best and Christiansen, 1991; Henry et al., 2011). The arrival of south-migrating ignimbrite flareup volcanism in the middle Cenozoic profoundly affected topography, disrupting hinterland drainage networks. This is most clearly shown by the systematic southward-younging ages of the oldest material recorded at the bases of west- and east-flowing Eocene–Oligocene paleovalleys (Fig. 2), suggesting that new drainages formed progressively southward, roughly synchronous with volcanism. In some areas, volcanism was also preceded by development of shallow basins, relatively minor offset along normal faults, and limited sedimentation. We suggest that volcanism caused pronounced uplift, perhaps of the order of 1.2 km based on the measured height of paleovalleys active during this time (Henry et al., 2012). Given the massive influx (see, e.g., Best et al., 2009) of heat associated with the addition of volcanic and plutonic material to the lithosphere, uplift is an expected result (e.g., Lachenbruch and Morgan, 1990), as exemplified today by the >1 km uplift around Yellowstone (e.g., Pierce et al., 1992). However, while the records of sedimentation, erosion, faulting, drainage development, and magmatism all support an elevation increase associated with the ignimbrite flareup, this is no longer clearly supported by the available stable isotope information from carbonates in basins across the western United States. A consequence of the improved age control and characterization of those sections near Elko (Lund Snee et al., 2016; Smith et al., 2017; this study) and in southwest Montana (Schwartz et al., 2019) is that prominent decreases in δ¹⁸O values during the early and middle Cenozoic (e.g., Horton et al., 2004; Davis et al., 2009b; Kent-Corson et al., 2006; Mix et al., 2011; Chamberlain et al., 2012) can no longer be tied directly to the onset of volcanism in these areas.

We propose that dynamic uplift accompanying ignimbrite flareup magmatism shifted the continental divide eastward into central Nevada from its prior position along the crest of the Cretaceous magmatic arc (Van Buer et al., 2009) and that this shift occurred in a southward-propagating fashion (Figs. 7 and 8). The middle Cenozoic highlands that supported this developing paleodivide were not static but instead responded dynamically as eruptions occurred and calderas formed. Volcanism left behind a plateau, with little documented erosion, tectonism, or sedimentation occurring until ca. 17 Ma (Figs. 6–8), although minor deposition initiated as early as latest Oligocene time in certain areas (Fig. 6), dominantly near developing metamorphic core complexes (Gans et al., 1989; Frerichs and Pekarek, 1994; Miller et al., 1999; McGrew and Snoke, 2015; Ruksznis, 2015; Lund Snee et al., 2016; this study). Elevations likely remained high following volcanism (perhaps with some component of gradual subsidence) due to the largely irreversible addition of magma to the crust and because rocks currently exposed in metamorphic core complexes experienced partial melting tens of millions of years later than the onset of Cenozoic magmatism (Howard et al., 2011; Strickland et al., 2011; Konstantinou et al., 2012, 2013a; Konstantinou and Miller, 2015). Rapid regional extension by Basin and Range faulting initiated ca. 17 Ma (Stockli, 2005; Colgan and Henry, 2009; Colgan, 2013), probably in response to changing boundary conditions, and driven by crust that remained elevated and thermally weakened.

The tectonic and topographic history and its temporal framework discussed here challenge suggestions that Mesozoic shortening produced a greatly thickened and elevated crust that drove gravitational collapse across the Sevier hinterland, either during shortening or soon after (see also Konstantinou, this volume). Evidence for rapidly evolving topography, drainage divides, and highlands related to and during Cenozoic magmatism also challenges the traditional notion of a long-lived, strongly elevated Nevadaplano (Cassel et al., 2012; Wells et al., 2012; Best et al., 2013) with a fixed (Late Cretaceous to) Cenozoic drainage divide (Henry et al., 2012). These suggestions pose implications for our understanding of orogenic and magmatic systems worldwide, underscoring the short time scales over which major changes in elevation and catchments can occur, particularly when the influence of magmatism on topography is considered.

We thank Trevor A. Dumitru for guidance with detrital zircon geochronology, Mark E. Raftrey for analyzing four detrital zircon samples, and Stephen Pearcey and Virginia Isava for assisting with sampling. The manuscript benefited from discussions with Simon L. Klemperer, George A. Thompson, Joseph P. Colgan, Karen I. Lund, Norm H. Sleep, and Theresa M. Schwartz. We are grateful to John P. Craddock for thoughtful editorial work and to Kathryn E. Snell, Theresa M. Schwartz, and David H. Malone for detailed comments that considerably improved the manuscript. This research was funded by National Science Foundation grant EAR-1322084 to Miller and a Stanford University G.J. Lieberman Fellowship to Lund Snee.