Stratigraphy of the Eocene–Oligocene Titus Canyon Formation, Death Valley, California (USA), and Eocene extensional tectonism in the Basin and Range

During the Cenozoic, the Basin and Range extensional province of western North America (Fig. 1) underwent distributed continental extension driven by a combination of external stresses (i.e., changing plate boundary stresses) and internal stresses (i.e., topographically induced gravitational potential energy gradients and buoyancy forces) (e.g., Sonder and Jones, 1999). Our ability to separate the relative influence of these geodynamic drivers relies on estimates of the initial state of topography and lithospheric structure prior to extension (e.g., Jones et al., 1996; Molnar et al., 2015; Levandowski et al., 2017; Ghosh et al., 2019; Zhou and Liu, 2019; Lund Snee and Zoback, 2020). Internal stresses are thought to have been a major driver of pre-Miocene extension in the Basin and Range, when significant gravitational potential energy gradients may have existed along the margins of the Nevadaplano, an inferred region of over-thickened crust constructed across western North America during the Cretaceous Sevier orogeny. External stress changes caused by plate boundary reorganization may also have played a role in continuing extension after initial orogenic collapse. Separating the roles of external stress changes and internal driving stresses on extension in the Basin and Range depends on knowing the location and rate of extension across the Basin and Range throughout the Cenozoic.

Early extension in the Basin and Range is thought to have started with a Cretaceous and Paleogene phase of relatively small magnitude synconvergent upper-crustal extension, followed by a later phase of larger-magnitude extensional collapse (Dickinson, 2002). Late Cretaceous and early Paleogene low-magnitude, syn-convergent extension is primarily inferred from a sparse stratigraphic record (e.g., Potter et al., 1995; Dubiel et al., 1996; Druschke et al., 2009; Lund Snee et al., 2016), and to a lesser degree from thermochronologic and geochronologic evidence (e.g., Wells et al., 1990; Applegate and Hodges, 1995; Camilleri and Chamberlain, 1997; Sullivan and Snoke, 2007). The record of Eocene to Miocene extension is both more widespread and indicative of larger magnitude extension that likely resulted in significant crustal thinning and topographic collapse, as recorded in thermochronologic data from metamorphic core complexes (e.g., Holm and Dokka, 1991; Lee, 1995; Foster and Raza, 2002; Vogl et al., 2012; Evans et al., 2015; Lee et al., 2017) and stratigraphic sequences deposited in hangingwall sedimentary basins (Satarugsa and Johnson, 2000; Wagner and Johnson, 2006; Smith et al., 2017; Canada et al., 2020).

The importance of internal stresses in localizing early extension in the Basin and Range is reasoned from the narrow north-south–oriented belt of Cordilleran metamorphic core complexes that lie to the west of the Sevier fold-thrust belt (Fig. 1; e.g., Coney and Harms, 1984). Localization of extension along this north-south belt is inferred to have been driven by enhanced topographic and buoyancy stresses in a continental lithosphere that was over-thickened by late Cretaceous shortening during the Sevier orogeny (Coney and Harms, 1984; Sonder et al., 1987; Jones et al., 1998). The presence of thickened crust in the Cordillera prior to extension is supported by multiple lines of evidence, including palinspastic reconstructions of extension (e.g., Coney and Harms, 1984; Bahadori et al., 2018), stable isotope proxy estimates of paleotopography (e.g., Chamberlain et al., 2012; Lechler et al., 2013; Cassel et al., 2014; Snell et al., 2014), and petrologic estimates of crustal thickness (Best et al., 2009; Chapman et al., 2015).

Despite compelling evidence of crustal thickening in the Sevier hinterland, the lag between the attainment of maximum crustal thickness and subsequent extensional collapse and crustal thinning (ca. 60 Ma; Chapman et al., 2015) suggests that the onset of extension may also require an external forcing. Two main mechanisms are proposed to have initiated early Basin and Range extension: the first is the timetransgressive, north-south removal of the Farallon slab and subsequent heating and weakening of the over-thickened crust (e.g., McQuarrie and Oskin, 2010); the second is a reduction in plate convergence rates resulting in plate boundary stress changes (e.g., Schellart et al., 2010). A particular challenge in disentangling the roles of these two external drivers is that a wealth of information regarding early Basin and Range extension is derived from studies in the northern Basin and Range (e.g., Potter et al., 1995; Druschke et al., 2009; Cassel et al., 2014; Lund Snee et al., 2016; Smith et al., 2017; Cassel et al., 2018; Long, 2018; Canada et al., 2019; Canada et al., 2020), where Eocene removal of the Farallon slab is contemporaneous with the ca. 40 Ma change in the rate of plate convergence and initiation of trench retreat along the plate boundary (Jurdy, 1984; Humphreys, 1995; McQuarrie and Oskin, 2010; Schellart et al., 2010). Farther south, in the central Basin and Range, Farallon slab removal occurred in the middle Miocene, at ca. 15 Ma, significantly later than the Eocene reduction in plate convergence, thus presenting a potential opportunity to examine the extensional response of the Basin and Range to slab removal and plate boundary reorganization separately. To leverage such an opportunity, however, requires the preservation of thermochronologic or stratigraphic evidence of Eocene extensional tectonism.

The Titus Canyon Formation, in eastern Death Valley (Fig. 1), is recognized as one of the earliest synextensional basin deposits in the central Basin and Range (Reynolds, 1969; Saylor, 1991; Snow and Lux, 1999; Fridrich and Thompson, 2011; Niemi, 2012). Previous detailed studies of the Titus Canyon Formation primarily focused on mapping its extent (Reynolds, 1969; Niemi, 2012) and describing its lithostratigraphy (Stock and Bode, 1935; Saylor, 1991; Gutenkunst, 2006), collectively providing limited geochronology and biostratigraphy to constrain the depositional age of the formation and thus the evolution of the Titus Canyon basin. We present ten new measured sections (Plate 1) and a detailed geologic map (Plate 2) of the Titus Canyon Formation in eastern Death Valley National Park. The stratigraphic sections and geologic mapping are accompanied by newly determined stratigraphic positions of paleontologic sites, detrital zircon U-Pb analyses used to inform both basin provenance and provide maximum depositional age (MDA) constraints, and compiled and recalculated ⁴⁰Ar/³⁹Ar geochronologic age data. We develop both an age model for the Titus Canyon Formation and a basin evolution model for the Titus Canyon basin in response to tectonic drivers. These models yield revisions to both the internal stratigraphy and stratigraphic continuity of the Titus Canyon Formation, as well as a reassessment of regional correlations of early Cenozoic stratigraphy in the Death Valley region. The age model and basin evolution model are considered in the context of Farallon slab removal and Eocene plate boundary reorganization to connect the development of the Titus Canyon basin to the initiation of extension across the entire Basin and Range province.

Paleogene strata in northern Death Valley that may record the earliest phases of extensional tectonism in the western Basin and Range are thin, discontinuous, and difficult to date. These challenges have inhibited the development of lithostratigraphic and chronostratigraphic models of these strata, resulting in a lack of consensus around regional correlation and temporal range of Paleogene rocks (Fig. 2; e.g., Snow and Lux, 1999; Fridrich and Thompson, 2011). We outline below several of the controversies surrounding these strata to illustrate the challenges associated with correlating strata across the highly extended Death Valley region.

Cenozoic strata in the Death Valley region were deposited in rapidly evolving extensional basins, resulting in significant lateral variability, time-transgressive deposition, and complex inter-play between deposition and active faulting (e.g., Çemen et al., 1999; Snow and Lux, 1999; Snyder and Hodges, 2000; Niemi, 2002; Fridrich and Thompson, 2011). Early efforts to map and define stratigraphic units within Death Valley were restricted to isolated study areas, resulting in nomenclature that varied from mountain range to mountain range for units that may in fact be lithologically or temporally correlative. Snow and Lux (1999) presented a tectono-stratigraphic framework that coherently placed these individual formations into allostratigraphic packages that they correlated across the entire Death Valley region, and in doing so defined three distinct episodes of basin development (Fig. 2). They termed these three phases the pre-extensional, synextensional, and post-extensional basin development stages, and although the phases are time-transgressive across the Death Valley region, the phases are identified by the development of regional unconformities, variations in depositional rates and styles, and variations in volcanic input. This tectono-stratigraphic framework was further expanded by Fridrich and Thompson (2011), who correlated seven tectono-stratigraphic assemblages across the Death Valley region (five of these seven assemblages that are geographically relevant to northern Death Valley are shown on Fig. 2). The oldest tectono-stratigraphic package, referred to as the Titus Canyon Assemblage (Fridrich and Thompson, 2011) or the Grapevine Sequence (Snow and Lux, 1999), contains two early-extensional units, the Titus Canyon Formation and the Ubehebe Formation; however, the depositional age, formational boundaries, and stratigraphic relationships between these two units remain unresolved (cf. Snow and Lux, 1999; Fridrich and Thompson, 2011).

The Titus Canyon Formation was first defined by Stock and Bode (1935), who were excavating fossil land mammal remains in the southern Grapevine Mountains (Fig. 1), including a Titanothere skull (Stock, 1936). Stock and Bode performed preliminary stratigraphic measurements and descriptions of the fossiliferous beds and defined the Titus Canyon Formation as an ~1000- m- thick sequence of brightly colored conglomerates, sandstones, limestones, and tuffs, but they did not provide detailed mapping of the areal extent of the unit or of its relationship to adjacent geologic units. The Titus Canyon Formation was originally assigned an Oligocene age based on the mammalian fossils found within it (Stock and Bode, 1935; Mason, 1988). Reynolds (1969) significantly refined the stratigraphy of the Titus Canyon Formation with the detailed geologic mapping in Titanothere and Titus canyons (Fig. 3) and also defined a type section for the formation in upper Titus Canyon (Fig. 3). Reynolds (1969) internally divided the Titus Canyon Formation into four units, from stratigraphically lowest to highest: (1) the sedimentary breccia facies; (2) the variegated facies; (3) the brown conglomerate facies; and (4) the green conglomerate facies, and provided an interpretation of the Titus Canyon basin as an extensional basin.

The stratigraphic definition of the Titus Canyon Formation by Reynolds (1969) was revisited with the expanded availability of geochronologic age control (Snow and Lux, 1999). In particular, ⁴⁰Ar/³⁹Ar ages bracketing an angular unconformity between the green conglomerate facies and the underlying variegated facies revealed a 15 m.y. hiatus in deposition, with an Oligocene depositional age for the variegated facies and a middle Miocene age for the green conglomerate facies (Saylor and Hodges, 1994; Snow and Lux, 1999). This discovery led to a revised stratigraphic definition of the Titus Canyon Formation that reassigned the green conglomerate facies to the Miocene Panuga Formation (Snow and Lux, 1999).

Geochronologic analyses throughout the Death Valley region led to the revised definition and regional correlation of multiple other early Cenozoic deposits. Early Cenozoic sedimentary strata from Bat Mountain (Fig. 1), in the southern Funeral Mountains (Fig. 1; McAllister, 1971; Çemen et al., 1982, 1985; Çemen and Wright, 1990) and the Cotton wood Mountains (Fig. 1; McAllister, 1952; Snow, 1990; Snow and White, 1990; Snow, 1993) were demonstrated to share latest Oligocene to early Miocene depositional ages, on the basis of intercalated tuffs (Çemen et al., 1999; Snow and Lux, 1999). These strata were defined as the Ubehebe Formation and consist of a fining-upwards sequence of conglomerates, sandstones, marls, and tuffs exposed in the Cottonwood and southern Funeral Mountains (Fig. 1; Çemen et al., 1982; Çemen et al., 1999; Snow and Lux, 1999). In both the Cottonwood Mountains and the Funeral Mountains (Fig. 1), the Ubehebe Formation is deposited directly on the sub-Tertiary unconformity (i.e., the regional unconformity across the Basin and Range defined by the deposition of Cenozoic strata on deformed Paleozoic or Mesozoic strata), and it is separated by a distinct angular unconformity from the overlying Panuga Formation (Snow and Lux, 1999).

In the southern Grapevine Mountains (Fig. 1), where the Titus Canyon Formation is defined, no strata demonstrably equivalent in age to the Ubehebe Formation have been previously recognized (Reynolds, 1969; Niemi, 2012). Snow and Lux (1999) considered the possibility that the Titus Canyon Formation might be an older, time-transgressive equivalent to the Ubehebe Formation, since Ubehebe Formation strata were not definitively observed stratigraphically above the Titus Canyon Formation anywhere, and the two units share some lithologic similarities. They concluded, however, that the Titus Canyon Formation is most likely distinct in age from, and thus not correlative with, the Ubehebe Formation. This conclusion hinged on reports that the Titus Canyon Formation is overlain by 20–24 Ma strata, equivalent in age to the Ubehebe Formation, at two isolated and relatively small, exposures (Reynolds, 1974; Wright and Troxel, 1993).

Strata broadly similar in lithology to the Titus Canyon and Ubehebe Formations have subsequently been reported along the eastern flank of Funeral Mountains (Fig. 1; Murray, 2002; Gutenkunst, 2006; Ridgway et al., 2011). Sections along the eastern Funeral Mountains were correlated by Gutenkunst (2006) to the Titus Canyon Formation on the basis of their similar lithofacies. However, geochronologic analyses from these sections ranged from latest Oligocene to early Miocene (Gutenkunst, 2006), which would suggest temporal equivalence to the Ubehebe Formation, not the Titus Canyon Formation. The continuum of exposures of early Cenozoic strata from the Grapevine Mountains southward along the eastern margin of the Funeral Mountains to Bat Mountain, at the southern end of the Funeral Mountains (Fig. 1), led Fridrich and Thompson (2011) to challenge the distinction of Snow and Lux (1999) and to propose that the Titus Canyon and Ubehebe Formations reflect a single, timetransgressive continuum of deposition. Support for this interpretation, in part, arises from a lack of definitive absolute age control that differentiates the depositional age of the Titus Canyon Formation from the Ubehebe Formation.

The challenge of resolving and understanding the early extensional history of the Death Valley region arises from two specific issues. First is the lack of absolute age control on the deposition of the Titus Canyon Formation, as discussed above. The age of the Titus Canyon Formation is primarily derived from paleontologic evidence and has been reported as Eocene to Oligocene (Stock and Bode, 1935; Mason, 1988; Lander, 2019), presenting difficulties in comparing the age of the Titus Canyon Formation to other lithologic units. The second issue is a lack of detailed mapping of the extent, continuity, and regional correlation of the Titus Canyon Formation, which had the added effect of preventing extrapolation of the limited biostratigraphic and geochronologic age constraints that do exist for the formation. Geologic maps at 1:48,000 scale in the Grapevine and Funeral Mountains (Cornwall and Kleinhampl, 1964; Wright and Troxel, 1993; Niemi, 2012) expanded the recognized areal extent of the Titus Canyon Formation beyond Titus Canyon to a northwest-southeast strip along the eastern Grapevine and Funeral Mountains (Fig. 3). However, these maps omit the internal stratigraphy of the Titus Canyon Formation as defined by Reynolds (1969), and, in areas particularly relevant to discriminating the Titus Canyon and Ubehebe Formations, predate the recognition of these two formations (e.g., Cornwall and Kleinhampl, 1964; Wright and Troxel, 1993).

Below, we present detailed measured sections (Plate 1) and geologic mapping (Plate 2) in the southern Grapevine Mountains and northern and central Funeral Mountains (Fig. 1) that inform the above-described controversies. The internal stratigraphy of the Titus Canyon Formation is mapped throughout the study region, with recognition of the possibility that the Ubehebe Formation can be differentiated from the Titus Canyon Formation. Robust geochronologic age constraints (Table 1) are provided for the type section of the Titus Canyon Formation and are integrated with previously unpublished age constraints from other sections to develop a coherent early Cenozoic chronostratigraphy for eastern Death Valley and resolve the age, stratigraphic definition of, and regional tectonic significance of, the Titus Canyon Formation.

To refine the internal stratigraphy and determine an age model for the Titus Canyon Formation, we measured stratigraphic sections, undertook large-scale field mapping, and completed zircon U-Pb geochronology. These efforts resulted in ten detailed measured stratigraphic sections (Plate 1), a 1:24,000 scale geologic map (Plate 2), and maximum depositional ages (MDAs) for the Titus Canyon Formation. These data were correlated with published and unpublished ⁴⁰Ar/³⁹Ar age determinations to develop a new basin evolution and regional stratigraphic model for the early-extensional sequences in Death Valley.

Geologic mapping included field and remote components. Fieldwork was conducted over two field seasons in 2019. We measured and described the internal stratigraphy of the Titus Canyon Formation using a Jacob’s staff at ten locations, comprising a total of 2600 m of strata. During geologic mapping in the field, the locations of faults and contacts were recorded directly onto high-resolution, ortho-rectified aerial imagery using a GPS-enabled tablet running Esri Collector for ArcGIS. Field mapping was principally focused on detailed analysis of Eocene to early Miocene stratigraphic units, with older and younger units mapped remotely (see index map on Plate 2). The locations and contacts of Proterozoic and Paleozoic units were principally derived from previous 1:48,000 and 1:100,000 geologic maps (Cornwall and Kleinhampl, 1964; Reynolds, 1969; Wright and Troxel, 1993; Carr et al., 1996; Niemi, 2012). Geologic contacts on those maps, which were originally on topographic base maps, were digitally refined to 1:24,000 scale by tracing color changes visible on high-resolution aerial imagery. In addition to ground truthing, some contacts between early Cenozoic sedimentary strata (including the Titus Canyon Formation) and overlying Miocene units were mapped by referencing the geologic mapping of Niemi (2012) and unpublished mapping by C. Fridrich (1999, personal commun.) and tracing contacts onto high-resolution aerial imagery. All field and remote mapping was compiled on a 1:24,000 U.S. Geological Survey topographic map base composed of digital line graphs (DLGs; 1980–1990 vintage), which were simplified for visual clarity (Plate 2). We precisely located Caltech and the Los Angeles County Museum paleontological localities excavated in the 1930s (Stock and Bode, 1935; Mason, 1988; Lander, 2019) by visually matching distinctive outcrops visible in original glass slide images of excavations to their modern exposures. We have purposefully obfuscated the exact geographic locations of fossil localities to protect the sites, although the stratigraphic position of such sites in reported measured sections (Plate 1) is representative of their observed positions within the Titus Canyon Formation.

New geochronologic analyses were undertaken by multiple methods; these new data were combined with previously unpublished ⁴⁰Ar/³⁹Ar ages to obtain an age model for deposition of the Titus Canyon Formation. All ages are summarized in Table 1.

Seven samples of tuffs and tuffaceous sandstones were collected from a range of stratigraphic levels within the Titus Canyon Formation for geochronologic analysis (Fig. 4 and Table 1). Selected samples appeared tuffaceous as assessed in the field with a hand lens. Zircon separates were produced at the University of Michigan following standard mineral separation techniques for rock crushing, sieving, and heavy liquids separation. U-Pb geochronology on detrital zircons (analytical data are available in Table S1¹) was performed using laser ablation–inductively coupled plasma mass spectrometry (LA-ICPMS) at the University of Arizona LaserChron Center following the analytical methods of Gehrels et al. (2008).

Maximum depositional ages were calculated from the zircon U-Pb ages for detrital zircon samples that contained Cenozoic zircons (four out of seven samples). There are a variety of published methods for MDA determination (e.g., Dickinson and Gehrels, 2009; Coutts et al., 2019), many of which are recognized to bias MDA estimates to young values when seeking the youngest zircon age or population (e.g., Coutts et al., 2019; Vermeesch, 2021). We present MDAs based on the Maximum Likelihood Age (MLA) algorithm, which is designed to minimize MDA bias (Galbraith and Laslett, 1993; Vermeesch, 2021). Application of the MLA method (Fig. 5) yielded consistent age determinations for a single tuffaceous unit dated by U-Pb detrital zircon MDA, ⁴⁰Ar/³⁹Ar, and zircon (U-Th)/He methods (Unit 38 tuff; Table 1), and consistent U-Pb detrital zircon MDA and ⁴⁰Ar/³⁹Ar ages determined on samples from two localities that were inferred to be the same geologic unit based on field mapping and stratigraphic relationships.

All Cenozoic zircon U-Pb ages from each detrital zircon sample were included in each MLA analysis, with the exception of a single 24 Ma zircon grain from the Unit 38 Tuff. This age was excluded because it could be demonstrated independently that the age disagreed with stratigraphic constraints on the depositional age of the unit sampled. Additionally, the measured U concentration from this grain (6675 ppm) was a factor of 6 higher than for any other grain analyzed, suggesting that the young age may be the result of diffusional lead loss due to radiation damage (Lee et al., 1997).

Four tuffaceous sandstones were dated with the (U-Th)/He method at the University of Michigan Thermochronology lab following the standard methods of Niemi and Clark (2018). Three grains were analyzed per sample and reported uncertainties are ± 1 standard error of the mean age of each sample. Ages are summarized in Table 1, and full analytical data are in Table S2 (see ^{footnote 1}).

No new ⁴⁰Ar/³⁹Ar analyses were performed for this study; however, six previously analyzed ⁴⁰Ar/³⁹Ar ages from tuffs within the Titus Canyon Formation and other Cenozoic strata in eastern Death Valley were recalculated and updated to the K decay constant of Renne et al. (2010, 2011) and the MMhb-1 monitor mineral age of Spell and McDougall (2003). Three unpublished ⁴⁰Ar/³⁹Ar ages from Cenozoic strata in the eastern Funeral Mountains (Gutenkunst, 2006) were recalculated from original analytical data using IsoplotR (Vermeesch, 2018) to determine plateau, isochron, weighted, and integrated ages (Fig. S1 and Table S3, see ^{footnote 1}). Two additional unpublished ⁴⁰Ar/³⁹Ar ages collected from the Titus Canyon Formation (Saylor and Hodges, 1994) were recalculated using the ArAR program (Mercer and Hodges, 2016). Original analytical data are not available for the samples from Saylor and Hodges (K. Hodges, personal commun., 2020). A single published ⁴⁰Ar/³⁹Ar age from the Panuga Formation in the southern Grapevine Mountains (Snow and Lux, 1999) was recalculated using ArAR (Mercer and Hodges, 2016). All ⁴⁰Ar/³⁹Ar ages are summarized in Table 1.

The accuracy of the geographic and stratigraphic positions of the ⁴⁰Ar/³⁹Ar samples is variable (location uncertainties are summarized in Table 1). Three sample locations from Gutenkunst (2006) were reported as stratigraphic position within a measured stratigraphic section from the eastern Funeral Mountains (section FM of Murray, 2002). The location of the FM measured section is approximated from a smallscale figure; so we employed field observations of the orientation and younging direction of strata in the field area of Murray (2002) to assess the location of the measured section and to estimate the geographic position of the samples in the field (Fig. 4). The two sample locations from Saylor and Hodges (1994) are unknown, but their approximate location within the stratigraphy was reported. An ⁴⁰Ar/³⁹Ar age from the Panuga Formation in the southern Grapevine Mountains (Fig. 1) is well defined stratigraphically (Snow and Lux, 1999). The written description of the sample locality within the Panuga Formation is consistent with the known distribution of this formation (Snow and Lux, 1999); however, reported geographic coordinates for this sample are inconsistent with both the written description and the distribution of the Panuga Formation. We therefore estimated the sample coordinates from the written description (Sample 592-GV1 K1; Fig. 4 and Table 1).

Our mapping supports the division of Paleogene strata in eastern Death Valley into the 40(?)– 30 Ma Titus Canyon Formation, and the ca. 28–19 Ma Rocks of Porter Mine, units that can be differentiated based on their lithologic content, differing volcanic content, distinct depositional timing based on radiometric ages, and consistent superposed stratigraphic relationship throughout the map area. Based on its distinctive lithology and our revised geochronology, the Rocks of Porter Mine may correlate with the Ubehebe Formation of Snow and Lux (1999). We describe both units below, with emphasis on the characteristics that distinguish them, both in the field and in terms of depositional age.

Our geologic mapping identifies exposures of the Titus Canyon Formation in a 35-km-long strip along the northeastern margin of the southern Grapevine and northern Funeral Mountains (Fig. 3 and Plate 2). The thickness of the formation varies from 500 m to ~1 km, with an overall trend of thickening to the northwest (Plate 2). Compared to previous studies (Stock and Bode, 1935; Cornwall and Kleinhampl, 1964; Wright and Troxel, 1993), our mapping extends the recognized distribution of the Titus Canyon Formation farther to the southeast. A small number of exposures originally mapped as the Titus Canyon Formation ~10 km northwest of the area mapped on Plate 2 (Stock and Bode, 1935; Saylor, 1991) have been previously refined as correlating with younger stratigraphic units (Niemi, 2012).

The depositional base of the Titus Canyon Formation is best exposed in the West Fork of Titus Canyon (in the vicinity of 36.8551°N, 117.0852°W), where the formation is deposited on the Cambrian Bonanza King Formation. Toward the southeast end of the mapping area near Indian Butte (Fig. 3, ~36.7015°N, 116.8117°W), the nature of the base of the Titus Canyon Formation becomes obscured by poor exposure and pervasive faulting but appears to lie primarily on the Cambrian Zabriskie Quartzite or Neoproterozoic to Cambrian Wood Canyon Formation (Plate 2). In many locations, especially near the southeastern extent of the map, extensive postdepositional normal faulting (Fig. 6) has put the Titus Canyon Formation in fault contact with underlying Paleozoic bedrock, making it difficult to determine the nature of the depositional base of the formation. Preserved exposures of the Titus Canyon Formation all lie in the hanging wall of the late Miocene Boundary Canyon detachment fault (Plate 2; Wright and Troxel, 1993).

Here, we redefine the internal stratigraphy of the Titus Canyon Formation from the previous descriptions provided by Reynolds (1969) and Snow and Lux (1999). Reynolds uses the term “facies” in the names of all his mapped sub-formation units (e.g., brown conglomerate facies). Our mapping indicates that the divisions between our proposed map units are laterally extensive and well defined; so we divide the stratigraphy of the Titus Canyon Formation into members, listed here from oldest to youngest: (1) the Basal breccia member (largely unmodified from Reynolds’ sedimentary breccia facies); (2) the Redbed member (rocks that were originally included in the lower part of the variegated facies of Reynolds); and (3) the Variegated member (the remainder of the variegated facies of Reynolds after removing the Redbed member). We include the brown conglomerate facies of Reynolds in the Variegated member.

The lowest unit of the Titus Canyon Formation, the Basal breccia member, includes rocks previously called the sedimentary breccia facies by Reynolds (1969). The member is relatively thin (~0–30 m) and discontinuous, even where it is best exposed in the west fork of Titus Canyon and along the Titus Canyon Road south-east of Red Pass (Fig. 3). The base of the member is marked by the first occurrence of angular fragments of Bonanza King Formation (Fig. 6B), which are typically deposited onto intact Bonanza King Formation. The member consists of massive or weakly stratified matrix-supported breccia and conglomerate with pebble- to boulder-sized (up to 2- m- diameter) clasts composed of primarily Paleozoic limestone, quartzite, and chert.

We define the Redbed member of the Titus Canyon Formation as a succession of recessive brickred calcareous siltstones to mudstones interfingered with resistant packages of grayishpurple sandstone and matrix-supported sandy conglomerate (Figs. 7B and 7D). This member approximately represents the lower third of the variegated facies defined by Reynolds (1969). The Redbed member comprises 250 m of strata at its thickest exposures, though many locations have as little as 50 m of section that may be truncated by faults (Plate 1). The Redbed member is present across the entire map area, identifiable by its distinct combination of purplishgray conglomerate and sandstone with brickred siltstone. The Redbed member fines upward overall, with its base consisting of predominantly purplish-gray cross-bedded sandy conglomerate that transitions upsection to recessive brick-red siltstone, which quickly become the dominant lithology toward the top of the member (Fig. 7A). While the transition from dominantly purple-gray conglomerate to brick-red siltstone occurs at ~150 m above the base of the member in the west fork of Titus Canyon (Fig. 3), elsewhere the transition varies in stratigraphic position within the member, and we therefore did not interpret the transition as a mappable unit contact. The red recessive exposures of the member are characterized by massive or weakly stratified brickred siltstone containing sparse secondary calcite veins, sparse light greenish-gray oxidation halos, and sparse knobby calcareous concretions. In some places, the recessive red siltstone is organized into 1–5-m-thick bedding packages that become slightly more resistant and calcareous up section before truncating at a planar surface and repeating (Fig. 7D). Intervals of resistant grayish-purple conglomerate contain well-rounded pebble- to cobble-sized clasts composed primarily of quartzite with some limestone and chert. Quartzite clasts are highly polished and have percussion marks. Conglomerate lenses generally fine upwards into grayish-purple sandstone consisting of fine to coarse subangular grains in planar beds or sets of 10–20-cm-tall crossbeds (Figs. 7C and 7D). The purple sandstone and conglomerate intervals weather into resistant, smooth, rounded outcrops, while the recessive brickred siltstone forms steep slopes mantled with loose weathered material (Figs. 7A and 7B). All the Duchesnean fossil localities (Stock and Bode, 1935; Lander, 2019) that we were able to locate lie within the brickred siltstone beds of the Redbed member (Plate 1).

We redefine the Variegated member of the Titus Canyon Formation from the description of Reynolds (1969) as a sequence of colorful siltstones, sandstones, conglomerates, and limestones containing three marker units, including a distinctive green sandstone and conglomerate referred to as the middle green sandstone, and two tuffs (units 8 and 38 of Reynolds, 1969; Plate 1). The Variegated member is no more than 500 m thick and comprises repeating sequences (10–20 m thick) of moderately resistant pale greenish- or yellowish-gray sandy conglomerate separated by 5–10-m-thick recessive packages of interbedded red calcareous siltstone and yellow limestone (Fig. 8A and Plate 1). The member outcrops as rounded hills that reflect differences in the erodibility of the conglomerate and siltstone intervals (Fig. 8A). The sandy conglomerate intervals (Fig. 8C) are typically matrix-supported, massive to cross-bedded, and contain 0.5–30-cm-diameter clasts of limestone, quartzite, and distinctive small black chert pebbles 0.5–3 cm in diameter. Conglomerates are separated by red calcareous siltstone intervals that contain common reduction halos (Fig. 8D) and 0.5–1.5-m-thick gray or red micritic limestones that in some places show thin bedding-parallel sparry laminae (Fig. 8E) and commonly weather to a distinctive pale yellow (Fig. 8E). The micritic limestone beds are mostly devoid of allochems with the exception of a few oncolite beds in Titanothere Canyon and near Indian Butte (Fig. 3); these beds contain abundant oncoids (Fig. 8G; Titanothere Canyon outcrop, 36.8298°N, 117.0128°W; Indian Butte outcrop, 36.7025°N, 116.8088°W). Limestone-rich intervals have been previously used as marker intervals in the Titus Canyon Formation, though their lateral continuity is not clear between different sections. The tracing of a single limestone bed over significant distances is made difficult by their relatively thin lenticular nature and the pervasive faulting in the area. As such, we cannot demonstrate the continuity of an “algal limestone” marker bed (Stock and Bode, 1935; Reynolds, 1969; Saylor, 1991) as a distinctive, traceable unit throughout the Titus Canyon Formation. Overall, the Variegated member does not display upsection changes in grain size and has a relatively consistent grainsize distribution. While the member is lithologically uniform through its entire thickness, with the exception of the Middle green sandstone marker unit, discussed below, sandy conglomerate intervals within the Variegated member do take on a redder oxide-stained appearance upsection (Plate 1).

The rhythmic bedding of the Variegated member is interrupted near its middle, where a distinctive sequence of greenish sandstones and conglomerates form a 50–60-m-thick marker unit originally identified by Stock and Bode (1935) as the “middle sandstone and conglomerate.” We dub this interval the middle green sandstone, a marker unit that can be found at most localities throughout the map area, and which exhibits somewhat variable lithology but always includes a massive, recessive bluish gray-green, well-sorted medium sandstone near its base; the sandstone weathers into distinctive spheroidal outcrops (Fig. 8F). Many exposures of this marker unit also include resistant greenish-gray, matrix-supported sandy conglomerate, and rarely recessive bluish green-gray mudstone and thin brown limestones. The conglomerates that accompany the middle green sandstone interval contain a similar clast composition to those in the rest of the Variegated member, but with a greener sandy matrix. The red siltstone intervals immediately above and below the middle green sandstone contain some of the most well-developed micrite limestone beds in the entire formation (Fig. 8B and Plate 1).

Within the Variegated member, we have mapped two additional marker beds previously identified by Reynolds (1969): the unit 8 tuff near the bottom of the member and the unit 38 tuff, which is near the top of the member. The unit 8 and unit 38 tuffs are tuffaceous sandstones with a slightly off-white color and fine powdery appearance. They appear rich in ash but in some places exhibit cross-bedding and a sandy composition that suggest they might be reworked, a possibility supported by the significant width of the Cenozoic age peaks in their detrital zircon spectra (Fig. 9).

We report new geochronologic age information for the Titus Canyon Formation and develop an age model for the formation, based on a variety of geochronologic approaches, and we discuss this age model with respect to recognized paleontologic age constraints.

Of four Titus Canyon Formation samples from the vicinity of Titus and Titanothere canyons in the southern Grapevine Mountains (Fig. 3) that contained Cenozoic zircons, three came from the Variegated member, and a fourth came from the top of the Redbed member (Fig. 9 and Table 1). Samples from stratigraphically lower in the Redbed member contained primarily Mesozoic zircons (Fig. 9). Maximum depositional ages, based on MLA analysis (Fig. 5), and reported in stratigraphic order, are 36.0 ± 0.2 Ma from near the top of the Redbed member, and 33.4 ± 0.2 Ma, 31.5 ± 0.8 Ma, and 30.6 ± 0.3 Ma from the Variegated member (Figs. 2 and 9; Table 1). Miller et al. (2022) reported six detrital zircon MDA determinations for samples collected from the Titus Canyon Formation near Leadfield and Monarch Canyon (Fig. 3); from oldest to youngest, these are 36.3 Ma, 36.0 Ma, 34 Ma, 33.0 Ma, 32.3 Ma, and 23.7 Ma (Table 1, Fig. 9).

Of the six published and unpublished ⁴⁰Ar/³⁹Ar ages previously determined on tuffs in the southern Grapevine and northern Funeral Mountains (Saylor and Hodges, 1994; Snow and Lux, 1999; Murray, 2002; Gutenkunst, 2006), two ages are from strata that we have mapped as the Variegated member of the Titus Canyon Formation (Table 1). One is a sample most likely collected from the unit 38 tuff, which yielded a sanidine ⁴⁰Ar/³⁹Ar age of 30.4 ± 0.6 Ma (Saylor and Hodges, 1994) consistent with the zircon MDA that we determined for the same unit (30.6 ± 0.3 Ma). The second is a sample from near the base of the Variegated member which yielded a biotite ⁴⁰Ar/³⁹Ar age of 34.5 ± 0.4 Ma (Saylor and Hodges, 1994). This age is similar to the 33.4 ± 0.2 Ma MDA from the base of the Variegated member near Leadfield (Fig. 3).

The zircon (U-Th)/He method can be used to determine eruptive ages of tuffs and other volcanic rocks (e.g., Niemi, 2012); however, due to the relatively low closure temperature of helium in zircon (~180 °C), this method is susceptible to resetting in cases where samples have been modestly buried or reheated. We measured four zircon (U-Th)/He ages on samples from the southern Grapevine Mountains (Fig. 4 and Table 1), three of which were collected from the Titus Canyon Formation in Titus Canyon (Fig. 3). Of these three samples, one from the unit 38 tuff near the top of the Variegated member yielded an age (28.9 ± 1.0 Ma) that overlapped within error with other radiometric age determinations for the unit 38 tuff. Two ages determined on the unit 8 tuff, near the base of the Variegated member, yielded ages younger than the stratigraphically higher unit 38 tuff, in one case by 15 m.y. We hypothesize that middle Miocene emplacement of at least six major ignimbrite sheets across the region (Plate 2; Niemi, 2012) may have variably reset underlying zircon (U-Th)/He ages in the Titus Canyon area by inducing hydrothermal circulation (e.g., Abbey et al., 2018). We therefore note when zircon (U-Th)/He ages agree with other geo- and thermochronometers, but we do not find that this method produces consistently reliable eruption ages in the Titus Canyon region.

Paleontologic evidence has been used to assign ages ranging from Eocene to Oligocene to the Titus Canyon Formation (Stock and Bode, 1935; Stock, 1936; Lander, 2019), with more recent studies coalescing around agreement on a Duchesnean North American Land Mammal age (middle to late Eocene, 37–40 Ma; e.g., Mason, 1988). A recent reassessment of fossils from the Titus Canyon Formation resulted in the identification of nine new local faunas, eight of which are found in the southern Grapevine Mountains in Titus and Titanothere canyons, all of which are inferred to be late Duchesnean in age (Lander, 2019).

The age of the base of the Titus Canyon Formation remains unresolved, as does the age of the Basal breccia member. The age of the Redbed member is partially constrained by both paleontologic and geochronologic data. Geochronologically, the 36.0 ± 0.2 Ma MDA from a sandstone near the top of the Redbed member, and the 34.5 ± 0.4 Ma ⁴⁰Ar/³⁹Ar from the base of the Variegated member tightly bracket the age of the top of the Redbed member at ca. 35 Ma. This age places the ca. 34 Ma Eocene–Oligocene boundary and climate transition (e.g., Liu et al., 2009; Fan et al., 2018) at or just above the Redbed–Variegated member contact. A Duchesnean age for fossils in the Titus Canyon Formation is consistent with the geochronologic constraints, as all the fossil localities are found within the Redbed member below the stratigraphic level of the 36.0 ± 0.2 Ma MDA constraint. We find the correlation between the Titus Canyon local fauna and the Upper Porvenir local fauna that yields an age of ca. 37–38 Ma at the stratigraphic position of the fossil localities within the Titus Canyon Formation to be reasonable, given other existing age constraints. In the western West Fork and Upper East Fork sections (Fig. 3 and Plate 1), 100–200 m of strata of the Redbed member underlie the stratigraphic position of the fossil localities. No Cenozoic zircons were recovered from samples collected from these strata. Given the ages of the fossils in the Titus Canyon Formation, and an estimated rate of deposition for the Redbed member (Fig. 9), it would be surprising if the base of the Redbed member were older than middle Eocene. However, determining the age of the lower portions of the Redbed member may require approaches with limited reliance on absolute age dating, such as magnetostratigraphy, and the coarse conglomeratic nature of these deposits may prove challenging in that regard.

The age of the Variegated member is well constrained by our data and spans from ca. 35 Ma at the base of the member, as described above, to ca. 30 Ma at the top, as recorded by multiple age measurements on the unit 38 tuff near the top of this member (Plate 1). No more than a few meters of Variegated member strata overlie the unit 38 tuff at any measured section, suggesting that this age is representative of the preserved top of the Titus Canyon Formation. Overall, the six MDAs from Miller et al. (2022) agree with our age model, with the lone exception of their sample ELM18DVTC-7, which produced a 23.7 Ma MDA despite lying distinctly below the confidently dated ca. 30 Ma unit 38 tuff (Fig. 9 and Plate 1). We observed no obvious structural or depositional structure that could explain such an age discrepancy, and further investigation will be necessary to reconcile these two ages.

In summary, the Titus Canyon Formation is middle Eocene to early Oligocene in age, contains a fossil record of Duchesnean land mammals, and encompasses the Eocene–Oligocene boundary.

In addition to their use in determining maximum depositional ages, detrital zircon U-Pb ages also provide information about provenance (e.g., Thomas, 2011). In samples from the Titus Canyon Formation, three distinct populations of zircon U-Pb ages are identifiable (Fig. 9). The oldest population of grains range from 3.1 Ga to 400 Ma, forming significant peaks at ca. 2.7 Ga, 1.9–1.8 Ga, 1.5–1.4 Ga, and 1.1–1.0 Ga (blue shaded peaks on Fig. 9). An intermediate age population forms a distinctive triple peak distribution of 105 Ma, 170 Ma, and 220 Ma (the prominent 170 Ma peak is shaded green on Fig. 9). The youngest zircon age population is Cenozoic and forms a cluster ranging from 40 Ma to 30 Ma (yellow shaded peak on Fig. 9). Broadly speaking, these three populations represent derivation from the Neoproterozoic and Paleozoic miogeocline (e.g., Gehrels et al., 1995), the Sierran arc (e.g., Lechler and Niemi, 2011), and Cenozoic volcanic centers of the Basin and Range (e.g., McQuarrie and Oskin, 2010).

The relative proportion of each of these three populations changes as a function of stratigraphic position in the Titus Canyon Formation (Fig. 9). The lowest three samples, all collected from the Redbed member, include almost entirely Mesozoic zircons, with no Cenozoic zircons and very few Paleozoic and Precambrian zircons. Samples from the Variegated member have a more variable detrital zircon age population distribution. Sierran zircons are the subordinate population in all samples from the Variegated member, while the major populations are characterized by either a Paleozoic and older miogeoclinal source or a Cenozoic volcanic source, presumably depending on whether the sample was dominated by fluvial erosion and transport or volcanic airfall deposition. The detrital zircon spectra of Miller et al. (2022) show a similar provenance pattern, with the upper half of the Titus Canyon Formation showing more Cenozoic volcanic zircons, fewer Mesozoic Sierran arc zircons, and more Paleozoic or older zircons than the lower half of the formation.

In the vicinity of the type section of the Titus Canyon Formation, near Leadfield (Fig. 3), the Variegated member of the Titus Canyon Formation is overlain in angular unconformity by distinctive green conglomerates of the middle Miocene Panuga Formation (Reynolds, 1969; Snow and Lux, 1999; Niemi, 2012). Both of these units can be traced southward along the eastern side of the Grapevine Mountains, where they remain juxtaposed across an unconformity. However, south of Daylight Pass (Fig. 3), which approximately separates the Grapevine Mountains from the Funeral Mountains, the two units continue to be distinguishable but are separated by an intervening package of strata that is not readily ascribed to either of these two units (Plate 2). Where exposures of this intervening sequence of strata have been previously mapped, they have been variably assigned to the Panuga Formation (or Green conglomerate facies) (Murray, 2002; Gutenkunst, 2006), the Titus Canyon Formation (Saylor, 1991; Gutenkunst, 2006) or referred to as Miocene Tuffaceous Sandstone and Volcanic Breccia (Wright and Troxel, 1993). We describe the lithostratigraphy and geochronologic age control on these strata below and demonstrate that these strata are lithologically and temporally distinct from the Titus Canyon Formation. We informally name these strata the Rocks of Porter Mine, for exposures south of Daylight Pass near Porter Mine (Fig. 3; 36.7741°N, 116.9129°W), and we discuss correlation of these strata with a succession of late Oligocene to early Miocene rocks that are regionally widespread in the Death Valley region and collectively termed the Ubehebe Formation (Snow and Lux, 1999; Fridrich and Thompson, 2011).

An intact section of the Rocks of Porter Mine previously mapped as Titus Canyon Formation (Saylor, 1991) is found in the Funeral Mountains, ~2 km southeast of Daylight Pass, ~1 km west of the abandoned Porter Mine (Fig. 3; Plate 2; 36.7741°N, 116.9129°W). These exposures can be traced south-eastward to exposures near Indian Butte (Fig. 3 and Plate 2; 36.6941°N, 116.8011°W) that were previously mapped as either Titus Canyon Formation or Panuga Formation (Saylor, 1991; Gutenkunst, 2006). The Rocks of Porter Mine consist of bluish-green and gray conglomerates with pebble- to cobble-sized clasts, coarse sandstones, and tuffaceous sandstones that fine upwards into yellowish-beige to white sandstones, brown resistant limestones, recessive pale mudstone, and thick airfall tuffs (Fig. 10). At the southeastern extent of the mapping area (Plate 2), the lowest part of the Rocks of Porter Mine consists of 1–4-m-thick greenish-gray sandy conglomerates interbedded with 1–4-m-thick beds of mudstone and coarse sandstones (Figs. 10D and 6G) that in some places take on a striking dark-green coloration. The unit fines upwards into less colorful beige, brown, and pale-yellow fissile shales, and 10- cm- to 1-m-thick fossiliferous limestones, some of which contain small gastropods (Figs. 10E and 10F). Fossil fish have also been reported from strata equivalent to those that we describe here (Nyborg et al., 2017).

Northwest of Daylight Pass in the Grapevine Mountains, exposures of the Rocks of Porter Mine are incomplete, with the best exposure being a fault-bounded wedge in upper Titanothere Canyon directly below Red Pass (Fig. 3; 36.8302°N, 117.0251°W). Here, sandstone, conglomerate, and mudstone beds (Fig. 10D) take on a light blue-green color, similar to the lowest part of the Indian Butte stratigraphy. Below Red Pass, the Rocks of Porter Mine also contain an undated, highly resistant welded tuff with pumice fragments (Figs. 10A, 10B, and 10C; 36.8299°N, 117.0259°W). Overall, the Rocks of Porter Mine thicken to the southeast, a trend opposite that of the underlying Titus Canyon Formation. The Rocks of Porter Mine also display a greater degree of lateral variability in thickness than the Titus Canyon Formation, possibly indicating deposition in small wedge-shaped basins bounded by syndepositional normal faults (e.g., Fridrich and Thompson, 2011).

Based on our field observations and mapping, the Rocks of Porter Mine include several ⁴⁰Ar/³⁹Ar age analyses that were previously reported as age constraints for the Titus Canyon or Panuga Formations but which instead represent the depositional age of the Rocks of Porter Mine (Table 1). Guten-kunst (2006) analyzed three ⁴⁰Ar/³⁹Ar ages from a measured section through the Rocks of Porter Mine near Indian Butte (Fig. 3 and Plate 2); we recalculated these ages to be 28.2 ± 0.2 Ma, 23.6 ± 0.1 Ma, and 19.3 ± 0.5 Ma (Fig. 4 and Table 1). A single sample from an exposure of the Rocks of Porter Mine near Chloride Cliff road (Fig. 3) produced a zircon (U-Th)/He age of 24.4 ± 2.1 Ma, which is consistent with ages from Gutenkunst (2006), but is subject to the caveats regarding potential resetting of zircon (U-Th)/He ages described in the Methods section (Table 1). In summary, recalculated ⁴⁰Ar/³⁹Ar ages from tuffs that we map as within the Rocks of Porter Mine span the interval 28.2 ± 0.2 Ma to 19.3 ± 0.5 Ma, indicating that these strata are approximately late Oligocene to early middle Miocene in age.

The late Oligocene to early middle Miocene depositional age range for the Rocks of Porter Mine is consistent with the depositional age range of other packages of strata previously correlated with the regionally widespread Ubehebe Formation. These include strata at Bat Mountain, at the southernmost end of the Funeral Mountains (25–20 Ma; Fridrich and Thompson, 2011; Fridrich et al., 2012) and in the Cottonwood Mountains (24–19 Ma; Snow and Lux, 1999). In addition to a similar age of deposition, the Rocks of Porter Mine share the fining-upward stratigraphy and volcanic-rich nature of other known exposures of the Ubehebe Formation, traits that differentiate this formation from the Titus Canyon Formation. Based on the depositional age determinations for the Rocks of Porter Mine, their distinct lithology, and similarities to other strata ascribed to the Ubehebe Formation, we tentatively correlate the Rocks of Porter Mine with the Ubehebe Formation of Snow and Lux (1999), pending identification and description of a type-section for this unit and targeted sampling for more comprehensive geochronologic constraints on the depositional age range of this unit.

The Titus Canyon Formation records deposition from the Duchesnean (37–40 Ma) to the early Oligocene (30 Ma) (Fig. 9). Our differentiation of the Titus Canyon Formation from the Rocks of Porter Mine affords us an opportunity to revisit the evolution of the Titus Canyon basin and argue that its extent is limited to exposures in the northern Funeral and southern Grapevine Mountains. Depositional patterns in continental extensional basins are primarily controlled by the structural production of accommodation space and the climatically modulated supply of sediment (Carroll and Bohacs, 1999; Gawthorpe and Leeder, 2000), with a balance between basins that overfilled with sediment and host throughgoing drainages (i.e., exorheic), or underfilled and hydrologically closed (i.e., endorheic). The relative proportion of fluvial to lacustrine strata in a basin can serve as a proxy for overfilled and underfilled conditions, respectively (e.g., Geurts et al., 2020), and provides a useful framework for dividing the depositional history of the Titus Canyon basin into two phases. The following interpretations build on extensive previous sedimentologic and stratigraphic work in this region (Stock and Bode, 1935; Reynolds, 1969; Saylor, 1991; Murray, 2002; Gutenkunst, 2006; Ridgway et al., 2011; Niemi, 2012; Miller et al., 2022). Our goal is to summarize and correlate these works and place their observations within a common tectonostratigraphic and geochronologic framework, as many of the previous studies were conducted in small, disconnected study areas without robust geochronology.

The Titus Canyon basin initiated as an incipient extensional basin no later than 37 Ma, and early deposition of the Basal breccia and Redbed members continued until ca. 35 Ma. The overall fining-upwards stratigraphy of the Basal breccia and Redbed members of the Titus Canyon Formation records a progression from proximal rock avalanche breccias and coarse fanglomerate deposits to distal overbank deposits and conglomeratic channel bodies (Reynolds, 1969; Murray, 2002; Gutenkunst, 2006; Ridgway et al., 2011). This progression is recognized in other early extensional basin sequences and arises from coarse scarpderived deposits eventually yielding to a basin-floor fluvial system as the basin area expands (Gawthorpe and Leeder, 2000). The dominance of fluvial over lacustrine deposition during this early phase of basin formation implies an overfilled basin condition, and thus an exorheic hydrology with throughgoing drainages. Estimates of depositional rates of ~25 m/m.y., as derived from the geochronology and biostratigraphy of the Red-bed member (Fig. 9) indicate a slower depositional rate than in the upper Titus Canyon Formation and a slower rate than observed in typical extensional basins (50–200 m/m.y.; Allen and Allen, 2013). Combined with the overfilled state of the basin, this implies a relatively slow tectonic subsidence rate during this period.

Provenance information from the Redbed member of the Titus Canyon Formation also sheds light on regional drainage geometries and extent at the time of Titus Canyon basin formation. Detrital zircon U-Pb provenance analysis of the Redbed member reveals consistent detrital zircon spectra throughout the unit, dominated by Mesozoic ages, with significant age peaks in the middle Jurassic (ca. 165 and 175 Ma), a broad peak of late Triassic ages, and subordinate Late Cretaceous ages (Fig. 9). These three age peaks are consistent with recognized pulses of Sierran arc magmatism (e.g., Ducea, 2001; Coleman et al., 2003; Cecil et al., 2019); however, a specific source for the observed age spectra in the Redbed member is challenging to pinpoint. The subordinate nature of the Late Cretaceous ages argues against derivation from the present-day eastern Sierra Nevada, where exposures of Late Cretaceous plutons are extensive (Chen and Moore, 1982; Coleman and Glazner, 1998).

The Titus Canyon basin was located just east of the boundary between primarily Mesozoic arc-related rocks of the Sierra Nevada to the west and Neoproterozoic through Paleozoic sedimentary strata of the Cordilleran miogeocline to the east (e.g., Jennings et al., 1977). Several ranges in the westernmost Basin and Range, including the White, Inyo, and Cottonwood mountains, expose minor Late Cretaceous and Triassic intrusive bodies and host large Jurassic batholiths, with ages that overlap the detrital zircon ages observed in the Redbed member (Chen and Moore, 1982; Snow et al., 1991; Niemi et al., 2001; Coleman et al., 2003), making them a viable source terrane for the early Titus Canyon Formation.

Conglomerate clast compositions in the Redbed member of the Titus Canyon Formation further support derivation from a westerly source (Reynolds, 1969). Although lower Paleozoic carbonate and siliciclastic strata are exposed in most of the ranges that surround Death Valley, the identification of case-hardened carbonate clasts as being derived from Pennsylvanian-aged limestone, and the identification of Permian fusulinid-bearing lime-stones (Reynolds, 1969), limit the source for the Redbed member conglomerate clasts to sources on the western side of modern-day Death Valley (McAllister, 1952; Niemi et al., 2001). Clasts of quartz monzonite of inferred Jurassic age, and porphyritic granite clasts, are also observed within the Red-bed member. The Cottonwood Mountains, on the northwestern side of Death Valley, are a compelling source terrane for the combination of Paleozoic miogeoclinal clasts, Jurassic-aged monzonite clasts, and upper Paleozoic fusulinid-bearing limestones (Snow and Lux, 1999; Niemi et al., 2001), as well as the Jurassic and Triassic detrital zircon spectra. Further geochronologic analysis of individual clasts is needed to link specific clasts to unique source terranes. However, despite uncertainty in the exact provenance of the Redbed member of the Titus Canyon Formation, the basin was receiving sediment from a catchment that encompassed rocks not exposed in the local bedrock of the Grapevine and Funeral mountains, and which was almost certainly derived from the west, an inference also supported by paleocurrent and sedimentary structure observations (Stock and Bode, 1935; Reynolds, 1969; Gutenkunst, 2006).

The Variegated member records a shift in lithologic character, provenance, and fossil content of the Titus Canyon basin. Within the Variegated, conglomeratic channel bodies and silty overbank deposits rhythmically alternate with 0.5–1.5-m-thick lacustrine micrite beds, contrasting with the coarse, proximal deposition recorded by the Redbed member (Reynolds, 1969; Murray, 2002; Gutenkunst, 2006; Ridgway et al., 2011). The appearance of lacustrine conditions above the Redbed–Variegated member contact implies a transition from entirely exorheic (fluvial) to sometimes under-filled lacustrine conditions that could have been tectonically or climactically modulated. However, the similar timing of the Redbed to Variegated member transition (approx. 36–34 Ma) to the Eocene–Oligocene climate transition (ca. 34 Ma), during which northern latitudes cooled ~5–7 °C, may support a strong climatic influence during this time (Liu et al., 2009; Fan et al., 2018). Potential climatic influence on the Titus Canyon basin is not addressed in this manuscript, but it remains an important area for future investigation.

Sediment accumulation rates and provenance data from the Variegated member, however, suggest an increase in tectonic activity within the Titus Canyon basin. Deposition rates for the Variegated member are as high as ~250 m/m.y., distinctly faster than the ~25 m/m.y. sediment accumulation rate within the Redbed member, and fast compared to compiled extensional basin rates of 50–200 m/m.y. (Fig. 9; Allen and Allen, 2013). Geologic evidence for increased rates of local tectonism as a cause for increased sediment accumulation rates are preserved within the Variegated member, which both coarsens and thickens into the Fall Canyon fault in the upper West Fork of Titus Canyon (Plate 2).

The increase in sediment accumulation rate in the Variegated member is accompanied by a shift in basin provenance signature. Detrital zircon spectra from the Variegated member record Cenozoic zircon grains, suggesting an increase in volcanic input to the basin, likely from the onset of Oligocene ignimbrite eruptions in northern and central Nevada (Coney, 1978; Best et al., 2013; Miller et al., 2022), although such a signature is not diagnostic of fluvial basin geometry or extent. More pronounced is the shift in provenance from a dominantly Mesozoic Sierran arc signature to a broad age spectrum, similar to the detrital zircon spectra of Paleozoic and Neoproterozoic miogeoclinal rocks. Three of the four detrital samples from the Variegated member yield no evidence of derivation from a Mesozoic arc source (Fig. 9), and in the one Variegated sample where Mesozoic ages are present, they are Late Jurassic to Early Cretaceous, markedly different from the mid-Jurassic and Triassic ages found in the Redbed member (Fig. 9). The greater variability in detrital zircon spectra observed within the Variegated member also reflects a departure from the Redbed member, which exhibits consistent detrital zircon signatures.

The provenance changes we observe in the detrital zircon age spectra between the Redbed and Variegated members match previous observations of changes in clast composition in the Titus Canyon Formation (Reynolds, 1969; Murray, 2002; Miller et al., 2022). The Variegated member does not contain clasts of upper Paleozoic (Permian and Pennsylvanian) carbonates, and granitoid clasts are less common than in the Redbed member (Reynolds, 1969). Within the Variegated member, Miller et al. (2022) identify several distinctive clast compositions that together suggest an origin from the Roberts Mountain and Golconda allochthons in northern Nevada. Miller et al. (2022) link red radiolarian chert pebbles and cobbles to the Upper Paleozoic Havallah sequence in the Golconda allochthon and cobbles of black chert with pale phosphatic nodules to Devonian deepwater strata of the Roberts Mountain allochthon. Miller et al. also identify cobbles of conglomerate which themselves contain clasts of mafic volcanics, radiolarian chert, and orthoquartzite that suggest derivation from conglomerates of the Antler basin of northern and central Nevada.

While the clast assemblage identified by Miller et al. (2022) points to a provenance from northern Nevada, a significant portion of Titus Canyon sediment must be derived from more local exposures of the Neoproterozoic and Paleozoic Cordilleran Miogeocline. Highly polished and well-rounded clasts within the Titus Canyon Formation comprise distinctive formations derived from the miogeocline, including the Stirling Quartzite, Wood Canyon Formation, Zabriskie Quartzite, and Eureka Quartzite (Reynolds, 1969; Niemi, 2002). The distribution of Paleozoic quartzites across the western U.S. makes these clasts of limited utility for ascribing provenance; however, the distribution of Neoproterozoic quartzites is limited to the western portion of the central Basin and Range (Stewart, 1970). In addition to highly polished and rounded pebbles and cobbles of quartzite and chert, the Redbed member also contains angular boulders of carbonate megabreccia (1−>10 m) that are derived from paleotopographic highs composed of local bedrock (predominantly the Paleozoic Bonanza King Formation), within the Titus Canyon Basin (Reynolds, 1969).

The well-rounded nature of the conglomerate clasts within the Variegated member, especially highly polished chert and quartzite clasts, along with the observation of clasts that are distally sourced from the Golconda and Roberts Mountain allochthons, have been used to argue that the Titus Canyon Formation represents a remnant of a larger fluvial system that transported sediment from sources in northern Nevada to the then coastal Mojave region (Miller et al. 2022). However, mature, polished, and well-rounded sedimentary clasts (especially quartzite and chert) can arise from multiple episodes of sedimentary recycling, as opposed to a singular episode of significant fluvial transport (e.g., Blatt, 1967; Horton and Schmitt, 1998). The observation that components of the Roberts Mountains/Golconda allochthon assemblage are themselves found in clasts of conglomerate within the Variegated member (Miller et al., 2022) provides further evidence of at least one generation of sedimentary recycling of the clast assemblage within the Titus Canyon Formation.

A potential source for recycled clasts is conglomerate of the Mississippian Eleana Formation, which is exposed immediately to the east of Titus Canyon at Bare Mountain and the Nevada Test Site (Fig. 1) and contains clasts of radiolarian chert, phosphatic chert, and mafic volcanics that are inferred to have been derived from erosion of the northern Antler foredeep during the Antler Orogeny (Trexler and Cashman, 1997). The Eleana Formation was deposited as a submarine fan in the southern Antler Basin foredeep (Poole and Sandberg, 1977) and contains quartzites associated with both allochthonous sources from northern Nevada and local (miogeoclinal) sources (Trexler and Cashman, 1997). Resistant clasts (cherts, quartzites, and litharenites) within the Eleana are wellrounded and polished (Trexler and Cashman, 1997). The Eleana Formation has been identified as a source of mafic volcanic and radiolarian chert clasts in other early Tertiary basins in the central Basin and Range (e.g., Goldstrand, 1992). As such, recycling of sediment from the Antler foredeep basin should be considered a possible proximal source for Titus Canyon Formation clasts that originated in the Roberts Mountains and Golconda allochthons (Miller et al., 2022).

The provenance of the Titus Canyon Formation has the potential to test hypotheses of the paleotopography and drainage network integration and reorganization of the Cordillera during the Eocene to Oligocene transition from contraction to extension, a topic of considerable debate (cf. Colgan and Henry, 2009; Henry and John, 2013; Snell et al., 2014; Smith et al., 2017; Cassel et al., 2018; Long, 2018; Best et al., 2016; Lund Snee and Miller, 2022; Miller et al., 2022). However, evidence of sedimentary recycling within the clast population of the Variegated member of the Titus Canyon Formation, and the possibility of a local source for the well-rounded and polished clasts of Roberts Mountains, Golconda, and miogeoclinal affinity contradict a straightforward interpretation of the provenance of the Titus Canyon Formation (e.g., Miller et al., 2022), and thus its implications for the paleogeography of the proto-Basin and Range.

The most compelling compositional change observed between the Redbed and Variegated members is the upsection loss of Sierranaffinity granitic clasts and miogeoclinal Permian and Pennsylvanian clasts. Together, these changes broadly suggest a shift from a westerly to a northerly sediment source, a shift also observed in our detrital zircon provenance data. The Late Jurassic to Early Cretaceous ages of detrital zircons in the Variegated member, which are ages not well represented among plutons of the Sierran arc, along with the lack of Late Triassic, Middle Jurassic, and Late Cretaceous zircon ages suggest that the source area of the Variegated member cannot contain significant exposure of the most common plutonic ages of the Sierran arc (e.g., Ducea, 2001; Cecil et al., 2019). The Sylvania and Palmetto mountains, which lie ~50 km north of presentday exposures of the Titus Canyon basin, are underlain by Late Jurassic plutons that are intruded by small plutonic bodies of Late Cretaceous age. These plutonic ages are consistent with the detrital zircon spectra in the Variegated member, although geochronologic age constraints on the crystallization ages of these plutons (as opposed to cooling ages) are sparse (e.g., Maldonado et al., 1988). We note that palinspastic reconstructions of the central Basin and Range place presentday exposures of the Eleana Formation (i.e., at Bare Mountain and in the Eleana Range) to the north of the Titus Canyon Basin in Eocene time (Snow and Wernicke, 2000; Cashman and Sturmer, 2021). Together, these observations delineate a proximal region to the north of the Titus Canyon Basin that would have been a viable source for the clast population and detrital zircon spectra observed in the Variegated member.

In summary, the Titus Canyon basin initiated as an extensional basin in the late Eocene on the eastern slope of the paleo–Sierra Nevada arc. Tectonic subsidence was limited in magnitude, as evidenced by the relatively small (~1 km, Plate 1) maximum thickness of the formation. Minimal subsidence would allow the basin to remain overfilled for most of its depositional history, although it does appear to have experienced transitions from endorheic to exorheic conditions in response to tectonic or climatic forcing, as is observed in other rift systems (e.g., the Rio Grande and East African rifts; Repasch et al., 2017; Berry et al., 2019). The consistency of the detrital zircon spectra and clast composition throughout the Redbed member suggest minimal drainage reorganization or evolution during this time. The Variegated member of the Titus Canyon basin records major changes in the Titus Canyon basin, including significantly increased sediment accumulation rates and a transition to a fluviallacustrine sedimentary environment, as well as a transition to a more northerly provenance source. Our interpretation of the basin as extensionally controlled is compatible with the basin being an intermediate sediment sink along a longer sediment transport path (e.g., Miller et al., 2022); however, as discussed above, it is not clear that available provenance data require distal sedimentary sources.

As one of the earliest recognized pre–Basin and Range extensional units in the Death Valley region (e.g., Stock and Bode, 1935; Reynolds, 1969), the Titus Canyon Formation has been tenuously correlated with multiple poorly dated units in the central Basin and Range. A number of presumed Paleogene rock exposures in and around the Death Valley region, including the Horse Spring Formation, the unnamed conglomerate of the Fallout and Jumbled Hills, the Rocks of Winapi Wash, and the unnamed conglomerate of Gravel Canyon, among others, have been correlated with the Titus Canyon Formation on the basis of broad lithologic similarity and similar stratigraphic position above the sub-Tertiary unconformity (Niemi, 2002, and references therein). However, these deposits have neither biostratigraphic nor absolute age control that substantiates a temporal correlation with the Titus Canyon Formation. Moreover, such lithostratigraphy-based correlations were largely proposed prior to the recognition of the complex early Cenozoic depositional history of the Death Valley region that has been delineated in the past decades (e.g., Çemen et al., 1999; Snow and Lux, 1999; Fridrich and Thompson, 2011), and that we further refine here. Descriptions of these units, gleaned primarily from stratigraphic descriptions on published geologic maps, generally bear greater similarity to the Rocks of Porter Mine (Ubehebe Formation) as defined and described here than to the Titus Canyon Formation. Frequent descriptions of fossiliferous limestones and well-developed tuff beds are particularly more diagnostic of Ubehebe-equivalent strata than of Titus Canyon. Thus, our inference is that the Titus Canyon basin was of limited spatial extent and is preserved almost exclusively within the southern Grapevine and northern Funeral Mountains, while the Ubehebe Formation may have a much broader spatial distribution across the central Basin and Range.

The Titus Canyon basin has been described as having an extensional or transtensional origin based on observations of basin-scale sedimentological architecture and facies analysis (e.g., Saylor, 1991; Gutenkunst, 2006; Fridrich and Thompson, 2011). However, a coherent structural framework for the Titus Canyon basin has not previously been defined, in part due to myriad geologic maps that partially encompass the extent of the Titus Canyon Formation (e.g., Cornwall and Kleinhampl, 1964; Reynolds, 1969; Saylor, 1991; Wright and Troxel, 1993; Niemi, 2012). In mapping the extent of the Titus Canyon Formation and synthesizing the structures that bound the formation, we find that the Titus Canyon Formation in the Grapevine Mountains is everywhere confined to the hanging wall of a complex fault system, the Fall Canyon–Titus Canyon fault system (Plate 2).

The Fall Canyon fault is a steeply eastdipping normal fault that extends northward from Titus Canyon, parallel to the range crest of the Grapevine Mountains (Plate 2). The Fall Canyon fault juxtaposes upright Paleozoic strata in the upper limb of a large west-vergent fold in its footwall (the Titus Canyon anticline–Corkscrew Peak syncline, or TCA-CPS) against Cenozoic strata, including the Titus Canyon Formation and younger strata, in its hanging wall (Figs. 11 and 12). Although the lower members of the Titus Canyon Formation are not exposed along the Fall Canyon fault, in the west fork of Titus Canyon (Fig. 3), conglomeratic intervals within the Variegated member both thicken and coarsen westward toward the Fall Canyon fault. No single conglomerate is thicker than ~40 m in our measured sections, but these conglomerates merge into a single, >100-m-thick package toward the Fall Canyon fault. The thickening and coarsening of the conglomerates may represent sedimentological variability within the basin (e.g., of alluvial fans), but we interpret them as growth strata associated with the Fall Canyon fault, which would suggest that the Fall Canyon fault may have been an Eocene–Oligocene basin bounding fault to the Titus Canyon basin. The extensional nature of the basin is also evidenced by the unconformable nature of the contact between the Titus Canyon Formation and overlying strata, along with the wedge-shaped map-view pattern of Titus Canyon Formation exposures, both of which would be expected with deposition into asymmetric fault-bounded half grabens (Gawthorpe and Leeder, 2000). These observations lead us to agree with previous interpretations of the Titus Canyon basin as an extensionally subsiding, asymmetric, and west-deepening basin (Saylor, 1991; Niemi, 2002; Gutenkunst, 2006; Fridrich and Thompson, 2011). Our extensional interpretation contrasts with other interpretations of the Titus Canyon basin as a tectonically quiescent, low-relief alluvial floodplain (Miller et al., 2022).

Where the trace of the Fall Canyon fault is exposed in Titus Canyon east of Leadfield (Fig. 3), it merges with the Titus Canyon fault, an enigmatic, nearly horizontal fault. The excellent exposure in Titus Canyon shows the Fall Canyon fault and Titus Canyon fault merging rather than crosscutting (Fridrich and Thompson, 2011). Additionally, the significant offset associated with the Fall Canyon fault does not continue into the footwall rocks of the Titus Canyon fault. The flat attitude of the Titus Canyon fault, and its exposure across the top of the southern Grapevine Mountains, originally led to the interpretation of this structure as a Mesozoic thrust fault (Reynolds, 1969). The recognition that this fault everywhere carries the depositional base of the Titus Canyon Formation, and that this fault is contiguous with the Fall Canyon fault, under-scores the nature of this structure as a low-angle Eocene detachment fault. The Titus Canyon fault places upright lower Paleozoic strata and overlying Titus Canyon Formation in its hanging wall on upper Neoproterozoic and lower Paleozoic strata of the overturned limb of the TCA-CPS (Fig. 12; Reynolds, 1969; Fall Canyon fault zone of Lutz et al., 2021). The Titus Canyon fault has multiple structural levels that appear to localize along relatively weak shale intervals within the Cambrian Carrara Formation and Neoproterozoic Wood Canyon Forma tion (Plate 2). The lower Paleozoic and upper Neoproterozoic strata are significantly structurally thinned (< 1 km thick) and form the floor of the Titus Canyon basin, on which the Basal breccia member of the Titus Canyon Formation was deposited. A top-to-the-southeast sense of displacement on the Titus Canyon fault is inferred from both its continuity with the east-dipping Fall Canyon fault and a likely northwesterly source for the upright Paleozoic strata carried in the Titus Canyon fault hanging wall (Niemi, 2012).

To the south, in the vicinity of Daylight Pass, all structures in the Grapevine Mountains (including the Titus Canyon fault and Titus Canyon anticline–Corkscrew Peak syncline) are truncated at the Boundary Canyon detachment (Fig. 11 and Plate 2), a top-to-the-northwest low-angle detachment fault that bounds the Funeral Mountains metamorphic core complex (Holm and Dokka, 1991; Hoisch and Simpson, 1993; Beyene, 2011; Fridrich and Thompson, 2011; Lutz et al., 2021). The Boundary Canyon detachment separates amphibolite-grade Neo-proterozoic sedimentary rocks in its footwall from zeolite-grade strata in its hanging wall (Hoisch and Simpson, 1993; Wright and Troxel, 1993; Mattinson et al., 2007; Verdel et al., 2011; Hoisch et al., 2014; Affinati et al., 2020). The metamorphic history of the footwall rocks exhumed by the Boundary Canyon detachment is complex (Hoisch and Simpson, 1993; Applegate and Hodges, 1995; Mattinson et al., 2007; Wells and Hoisch, 2008; Affinati et al., 2020), but the timing of metamorphic core complex development is well constrained to the middle to late Miocene by a variety of low-temperature thermo-chronometric studies (e.g., Holm and Dokka, 1991; Lutz et al., 2021).

Where the Boundary Canyon detachment truncates the Titus Canyon fault along the eastern margin of the northern Funeral Mountains (near Indian Butte, Fig. 3), exposures of the Titus Canyon Formation are incomplete and primarily consist of the Variegated member of the Titus Canyon Formation (Plate 2). Near the Boundary Canyon detachment, the base of the Titus Canyon Formation is commonly faulted and juxtaposed against upper Neoproterozoic siliciclastic strata in the footwall of the detachment, recording disarticulation of the Titus Canyon Formation and associated Eocene structures during Miocene displacement on the Boundary Canyon detachment (Plate 2; Wright and Troxel, 1993; Gutenkunst, 2006).

Evidence for Miocene reactivation of the Eocene–Oligocene structures that bound the Titus Canyon basin is equivocal. Thick sequences of middle Miocene strata and ignimbrites from the Southwest Nevada volcanic field are exposed to the east of the Fall Canyon fault but are conspicuously absent to the west of this structure, throughout the western central Basin and Range (Niemi, 2012). Despite the thickness of these Miocene volcanic and sedimentary units, there is no clear evidence of thickening or growth fault relationships between them and the Fall Canyon fault. Moreover, we observed no unambiguous exposures that would clarify the nature of the contact between the Miocene units and the Fall Canyon fault as being structural rather than, for example, a depositional contact along a buttress unconformity. Given a lack of conclusive evidence for post-Oligocene activity on the Fall Canyon and Titus Canyon fault, we infer that the relationship observed between the Miocene ignimbrites and the Fall Canyon fault represents a buttress unconformity, but we acknowledge the challenge of preserving Oligocene structural relief into the middle Miocene.

Palinspastic restorations are key to understanding the magnitude and mechanisms of largescale intracontinental extension (e.g., Wernicke et al., 1988; Snow and Wernicke, 2000; McQuarrie and Wernicke, 2005), and they form the basis for geo-dynamic models constructed to understand the internal and external forces that drive lithospheric deformation (e.g., Bahadori and Holt, 2019). In the central Basin and Range, geologic markers that constrain such reconstructions fall into two broad categories: (1) late Paleozoic to Mesozoic contractional structures formed during convergence on the North American plate boundary (e.g., Burchfiel and Davis, 1975; Giallorenzo et al., 2018), and proximal sedimentary deposits of late Cenozoic age that have been tectonically dismembered by extension (e.g., Salyards and Shoemaker, 1987; Topping, 1993; Niemi et al., 2001). The structural markers are commonly used to define the magnitude of Basin and Range extension but have limited utility for understanding timing or rates of extension, while synextensional sedimentary deposits can provide compelling information on the timing of deformation but generally less precise constraints on extensional magnitudes.

The Titus Canyon Formation and bounding structures are thus of significant interest for refining palinspastic reconstructions because both closely pre-date the onset of large-magnitude extension in the Basin and Range. In this context, the structural relationship between the Fall Canyon–Titus Canyon fault system and the Titus Canyon anticline–Corkscrew Peak syncline is of particular significance. The TCA-CPS is a west-vergent fold pair, a rare type of contractional structure in the Death Valley region, which is dominated by east-directed thrust faults (e.g., Snow and Wernicke, 1989). As such, the TCA-CPS has been correlated with west-vergent folds in other mountain ranges surrounding Death Valley as a diagnostic structural marker (e.g., Wernicke et al., 1988; Snow and Wernicke, 1989), a correlation predicated on the assumption that the TCA-CPS, as well as the other west-vergent folds, are extensionally disrupted fragments of a single, originally contiguous geologic structure.

The age of west-vergent folding in the Death Valley region is constrained only in the Cottonwood Mountains, where crosscutting dikes and growth strata bracket the timing of folding to Late Permian to early Triassic age (Snow et al., 1991; Stevens and Stone, 2005). Consequently, the question of whether all west-vergent folds throughout the Death Valley region are of this age, and thus correlative, is debated (e.g., Serpa and Pavlis, 1996; Renik and Christie-Blick, 2013; Lutz et al., 2021). In particular, the TCA-CPS has been the subject of debate with respect to the timing of its formation, with similarities to structures in the Cottonwood Mountains supporting a Permo-Triassic age (Snow and Wernicke, 2000; Niemi, 2012). Alternatively, the close association of the TCA-CPS to the late Miocene Boundary Canyon detachment has been used to argue for a model of synextensional folding synchronous with unroofing of the Funeral Mountains metamorphic core complex (Reynolds, 1974; Lutz et al., 2021).

Age constraints on the Fall Canyon fault–Titus Canyon fault, and crosscutting relationships between that structure and the TCA-CPS, do not provide absolute age constraints on the timing of TCA-CPS folding, but do provide relative timing on the development of these structures. The normal fault emplacement of upright Paleozoic strata and Titus Canyon Formation onto overturned Neoproterozoic strata of the Corkscrew Peak syncline (Fig. 12) requires that the TCA-CPS was developed prior to displacement on the Fall Canyon–Titus Canyon fault. The coarsening and thickening of the Titus Canyon Formation into the Fall Canyon fault in western Titus Canyon therefore imply that the TCA-CPS formed prior to the Eocene. This age constraint for folding of the TCA-CPS does not resolve all outstanding questions of structural correlations between west-vergent folds in the Death Valley region but does support the inference that the TCA-CPS is a pre-Cenozoic contractional structure associated with late Paleozoic or Mesozoic shortening of the Cordilleran margin.

A variety of geodynamic mechanisms have been invoked as proximate causes for the onset of Eocene extension within the proto-Basin and Range, including flexural and isostatic dynamics ahead of the delaminating Farallon slab (e.g., Smith et al., 2017; Cassel et al., 2018), changing plate boundary stresses (e.g., Atwater, 1970; Schellart et al., 2010), and gravitational collapse (Humphreys, 1995; Sonder and Jones, 1999). In northern Nevada, the close temporal association between Farallon slab removal and the onset of extension has been presumed to reflect a causal relationship between the two (e.g., Cassel et al., 2018). However, farther to the north, the earlier removal of the Farallon slab has been linked to an early Eocene phase of volcanism and extensional deformation that preceded a late Eocene phase of deformation, potentially related to a change in plate boundary stresses (e.g., Janecke, 1992). The southerly location of the Titus Canyon Basin may potentially record the opposite scenario, in which a late Eocene change in plate boundary stresses preceded the removal of the Farallon slab by 15 Ma, providing a unique and important constraint on the evolution of internal and external forces driving Basin and Range extension.

In the central and northern Basin and Range, Late Cretaceous to Paleocene syn-convergent extension (e.g., Hodges and Walker, 1992; Wells et al., 2008; Druschke et al., 2009; Long et al., 2015) left a relatively sparse stratigraphic record in the hinterland of the Sevier orogen, primarily consisting of the Sheep Pass Formation (ca. 70–43 Ma) in central Nevada (Fig. 13, Druschke et al., 2009, 2011). Despite a sparse surficial record of syn-convergent extension, thermochronometric and isotopic data sets from several metamorphic core complexes record cooling during the late Cretaceous to early Cenozoic that may be attributable to early extensional exhumation, including in the Raft River–Albion metamorphic core complex (Wells et al., 2012), the Funeral Mountain metamorphic core complex (Applegate et al., 1992; Hoisch and Simpson, 1993; Applegate and Hodges, 1995), the Eureka Culmination (Long et al., 2015), and the Pequop Mountains (Camilleri, 1996). Syn-convergent extension has been attributed to collapse driven by gravitational potential energy gradients produced by a lithosphere over-thickened by Sevier deformation (e.g., Vandervoort and Schmitt, 1990; Hodges and Walker, 1992; DeCelles, 2004; Platt, 2007; Druschke et al., 2009).

Compared to earlier deposition, the Eocene stratigraphic record of extensional tectonism is more widespread and better constrained temporally (Fig. 13). This stratigraphic record preserves a north-to-south sweep in the onset of Eocene basin deposition, with the earliest deposits preserved in the Challis volcanic field of southern Idaho, where a basal conglomerate records deposition prior to the eruption of overlying volcanic units that have been dated with ⁴⁰Ar/³⁹Ar at ca. 49.5 Ma (Fig. 13; Moye et al., 1988; Sanford, 2005). In northernmost Nevada, the Elko Formation (Elko Basin) and the Dead Horse Formation (Copper Basin) contain a large number of tuffs that have enabled robust ⁴⁰Ar/³⁹Ar geochronology that indicates deposition spanned ca. 46–38 Ma, with the exception of one earlier local pulse of deposition in the vicinity of Coal Mine Canyon at ca. 49 Ma (Rahl et al., 2002; Haynes, 2003; McGrew et al., 2007; Henry, 2008; Henry et al., 2015; Mulch et al., 2015; Lund Snee et al., 2016; Smith et al., 2017). The synconvergent Sheep Pass Formation in central Nevada is overlain by Eocene strata, including the Stinking Springs Conglomerate and Kinsey Canyon Formation with depositional ages that range from ca. 37–35 Ma (Fig. 13; Druschke et al., 2009). Eocene deposition in the Sheep Pass Basin, in particular, overlaps with Redbed member deposition in the Titus Canyon basin.

Each basin displays a broadly similar depositional pattern, with early deposition dominated by coarse basal deposits and fluvial strata that transition up section into lacustrine strata capped by volcanic or volcanicrich units (Fig. 13). In the northern Basin and Range, the close temporal connection between clastic basin deposition and overlying volcanic deposition, combined with the southward sweep of basin initiation, has been taken to reflect subsidence and volcanism arising from the southward-propagating removal of the Farallon slab (e.g., Humphreys, 1995). In such a model, north to south delamination of the Farallon flat slab produced a south-directed, time-transgressive sweep of volcanism (Fig. 14; Humphreys, 1995; McQuarrie and Oskin, 2010), core complex exhumation (Fig. 14; e.g., MacCready and Snoke, 1997; Methner et al., 2015; Lee et al., 2017), and hinterland basin deposition (Fig. 14; Smith et al., 2017; Cassel et al., 2018; Lund Snee, 2020).

The southerly location of the Titus Canyon basin with respect to the time-transgressive removal of the Farallon slab, however, yields a markedly different pattern (Fig. 13). Eocene basin initiation and clastic deposition occurred in the Titus Canyon basin coeval with Eocene basin formation in northern Nevada, but the detaching Farallon slab, and associated volcanism, did not reach the central Basin and Range until the middle Miocene, coincident with widespread ignimbrite production in the Southwest Nevada volcanic field (Sawyer et al., 1994) and the exhumation of the Funeral Mountains metamorphic core complex (Holm and Dokka, 1991; Hoisch and Simpson, 1993; Lutz et al., 2021).

In the northern Basin and Range, Eocene hinterland basin development could reasonably represent a dynamic response to the southward removal of the Farallon slab (Smith et al., 2017; Cassel et al., 2018; Lund Snee, 2020), but the 15 m.y. gap between Eocene basin formation and volcanism and metamorphic core complex formation in the central Basin and Range would appear to be inconsistent with a singular driving force at the latitude of Titus Canyon. Throughout the Paleogene, ongoing subduction resulted in the fragmentation of the Farallon slab (e.g., Schellart et al., 2010). Between ca. 48 and 30 Ma, the trench-parallel length of the Farallon slab decreased by 75% and the Farallon–North America convergence rate rapidly slowed (Fig. 13; Jurdy, 1984; Gordon and Jurdy, 1986; Schellart et al., 2010). From the Eocene to Miocene, the Farallon slab steepened or foundered, from north to south across the northern Basin and Range (Fig. 13; Humphreys, 1995; McQuarrie and Oskin, 2010). The slowdown in Farallon–North America plate convergence rate was less time transgressive than Farallon detachment, consisting of a rapid transition from trench advance to trench retreat along the full length of the Farallon plate boundary at ca. 40 Ma (Fig. 13; Schellart et al., 2010).

In the northern Basin and Range, the transition from trench advance to trench retreat along the Farallon–North America plate boundary is temporally indistinguishable from the passage of the detaching Farallon slab (Fig. 13). However, these two events are separated by ~25 m.y. at the latitude of the Titus Canyon basin. Extensional basin formation and deposition of the Titus Canyon Formation is approximately coeval with the transition from trench advance to trench retreat along the plate margin, while the subsequent onset of local volcanism and metamorphic core complex formation is contemporaneous with the removal of the Farallon slab at that latitude (Fig. 13). We infer that the complex history of extensional tectonism, deposition, and basin formation in the central Basin and Range results from multiple discrete episodes of deformation. In the Eocene, extensional basin formation was driven by stress changes along the plate boundary, coupled with localized gravitational collapse of possibly over-thickened lithosphere (e.g., Vandervoort and Schmitt, 1990; Hodges and Walker, 1992; Platt, 2007; Druschke et al., 2009; Wells et al., 2012; Bahadori et al., 2018). These early extensional basins were overprinted by middle Miocene volcanism and core complex formation that accompanied extension driven by lithospheric weakening and heating during removal of the Farallon slab (e.g., Armstrong and Ward, 1991; Axen et al., 1993; Humphreys et al., 2003). Such a model is compatible with early clastic basin formation and subsequent volcanism and core complex formation in the northern Basin and Range, but the temporal overlap between these two geodynamic processes precludes their effects from being assessed separately in that region.

Thickened crust within the Sevier orogenic belt is hypothesized to have supported a highelevation, low-relief orogenic plateau called the Nevadaplano (Ernst, 2009). The spatial extent and elevation of this plateau are the subjects of considerable debate (e.g., Chamberlain et al., 2012; Long, 2012; Lechler et al., 2013; Snell et al., 2014; Miller et al., 2022) and are of particular interest in relating the distribution and magnitude of Basin and Range extension to the pre-extensional distribution of topography and crustal thickness (e.g., Bahadori et al., 2018). Approaches that directly constrain pre-extensional topography of the Nevadaplano are limited by the small number of deposits of appropriate age and composition for paleoelevation analysis, the thermal resetting of paleoaltimetric proxies such as carbonate clumped isotopes (Lechler et al., 2013), and the relatively large elevation uncertainties associated with paleofloral data and stable isotopic proxies (e.g., Chamberlain et al., 2012; Snell et al., 2014). The low spatial and temporal density of paleoelevation, paleotopographic, and paleocrustal thickness estimates has resulted in a number of competing models for the topographic profile of the Sevier orogenic belt (Cecil et al., 2010; Cassel et al., 2014; Bahadori and Holt, 2019). Among the most pronounced discrepancies in these models is the recognition of a paleotopographic divide, based on ignimbrite sheet distributions (Henry et al., 2012; Best et al., 2016), that lies ~150 km west (in pre-extensional coordinates) of the belt of core complexes and presumed greatest pre-extensional crustal thickness (Fig. 14; Bahadori and Holt, 2019).

The distribution of Eocene hinterland extensional basin deposits, including the Titus Canyon Formation, sheds some light on the potential paleotopography in the hinterland of the Sevier orogenic belt. All Eocene basins lie east of the paleotopographic divide and west of the inferred axis of thick crust produced by Sevier shortening (Fig. 14). In particular, the large areal extents of the lacustrine-dominated Elko Formation and the Sheep Pass Formation imply basin development across a region of relatively low-relief topography. An interpretation of low-relief paleotopography in this region is supported by the general lack of structural relief across the inferred Neva-daplano (Gans and Miller, 1983; Miller and Gans, 1989; DeCelles and Coogan, 2006; Van Buer et al., 2009; Konstantinou et al., 2012; Long, 2012, 2018), and the local relief on the base of ignimbrite sheets, as well as their long outflow distances and broad aprons (Henry et al., 2012). Together, these observations suggest the existence of a low-relief, and perhaps at times internally drained, portion of the Nevadaplano that can be extended along the belt of Eocene basins, from the Copper Basin in northern Nevada to the Titus Canyon basin in eastern California (Fig. 14). The extent of large paleolakes and lacustrine basins has been invoked as a proxy for low-relief topography on elevated plateau surfaces (Zhenhan et al., 2008), and the presence of such basins along the axis of the Sevier orogenic belt in Eocene time should be considered in paleotopographic reconstructions of western U.S. topography at that time (e.g., Bahadori and Holt, 2019).

In this study, we redefine the internal stratigraphy of the Eocene–Oligocene Titus Canyon Formation and clarify its correlation to regionally similar units through a combination of detailed stratigraphic descriptions and geochronology. A combination of mammalian biostratigraphy, new detrital zircon U-Pb maximum depositional age constraints, and relocated ⁴⁰Ar/³⁹Ar ages indicate that the Titus Canyon basin spans ca. 40–30 Ma and may record the earliest hinterland sedimentation in the central Basin and Range following the Mesozoic Sevier orogeny. The stratigraphy and structural position of the Titus Canyon Formation in the hanging wall of the linked Fall Canyon and Titus Canyon faults support previous interpretations of the formation as an extensional basin deposit. A change in detrital zircon provenance from an entirely Mesozoic Sierran arc source to a dominantly Paleozoic and Neoproterozoic miogeoclinal source may record progressive extensional unroofing of deeper miogeoclinal formations that contain greater concentrations of zircons, which along with the basin subsidence itself, supports continued extensional tectonism throughout the deposition of the formation. The phase of low-magnitude extension recorded by the Titus Canyon Formation predates large-magnitude exhumation of the nearby Funeral Mountains metamorphic core complex and suggests that extensional deformation in the central Basin and Range may have initiated because of changes in plate boundary forces and was later over-printed by deformation associated with removal of the Farallon slab.

This work was supported by the U. S. Geological Survey National Cooperative Geologic Mapping Program under USGS award number G19AC00138, with matching support from the University of Michigan. The views and conclusions contained in this document are those of the authors and should not be interpreted as necessarily representing the official policies, either expressed or implied, of the U.S. Government. Additional support came from National Science Foundation grant EAR-1151247 to NAN, a Geological Society of America Graduate Student Research Grant to NM, and a University of Michigan Department of Earth and Environmental Sciences Turner Award to NM. Fieldwork in Death Valley National Park was conducted under permit DEVA-2017-SCI-0037, issued to Torrey Nyborg. John Stark (DVNP) provided access to National Park Service GIS data. We thank Bruce Lander for providing the impetus to revisit the Titus Canyon Formation, and for many stimulating conversations with both him and Torrey Nyborg in the field. Tom Kelly, Brandt Nyborg, and John Sifling are thanked for participating in field expeditions. Thank you to Ken Ridgway for granting permission to publish ⁴⁰Ar/³⁹Ar data produced under U.S. Nuclear Regulatory Commission Contract NRC-02-07-006. This manuscript was improved by constructive criticism from Marin Clark, Michael Wells, Jens-Erik Lund Snee, Ryan Leary, and Jeff Lee. We acknowledge the Indigenous peoples who are the original inhabitants of the lands this work examines—the Newe Segobia (Western Shoshone) and Nuwuvi (Southern Paiute) tribes. We thank Mel for breakfast.