Progressive Miocene unroofing of the Big Maria and Riverside Mountains (southeastern California, USA) along the southwestern margin of the Colorado River extensional corridor

The Colorado River extensional corridor (CREC) in southwestern North America represents a highly extended region of continental crust in the Basin and Range (e.g., Crittenden, 1980; Rehrig and Reynolds, 1980; Lister and Davis, 1989). Ductile shear zones and brittle low-angle detachment faults associated with the metamorphic core complexes of the CREC facilitated large displacements (more than tens of kilometers) and exhumed middle- and upper-crustal rocks in their footwalls. These meta morphic core complexes characterize the interior portion of the Cordilleran orogenic belt of western North America, spanning from Sonora (Mexico) to British Columbia (Canada), and were critical in the extensional dismemberment of the Jurassic–Cretaceous Cordilleran contractional hinterland (e.g., Coney, 1980; Crittenden, 1980; Armstrong, 1982). Considering their tectonic importance in structurally accommodating large-magnitude extension in retroarc and backarc regions; reconstructing the geometric, temporal, and thermal evolution of detachment faults; and understanding the role of crustal and structural inheritance are critical.

Mountain ranges in the CREC have been the sites of numerous key studies that have helped formulate fundamental concepts and models on the structural evolution of detachment faults and corresponding metamorphic core complexes (e.g., Rehrig and Reynolds, 1980; Lister and Davis, 1989; Howard and John, 1987; Spencer and Reynolds, 1989, 1991; John and Foster, 1993). This foundational understanding and recognition of low-angle normal faults developed in the CREC has since been applied worldwide, such as in Papua New Guinea (e.g., Österle et al., 2021), the Greek Cyclades (e.g., Lister et al., 1984; Brichau et al., 2006; Grasemann et al., 2012), and the North China craton (e.g., Yang et al., 2007).

Low-temperature thermochronometry is a useful technique for quantifying the timing and magnitude of crustal exhumation in different tectonic regimes, including contractional, extensional, and transpressional- to- transtensional systems (e.g., Fitzgerald et al., 1991; Foster and John, 1999; Ehlers et al., 2005; Stockli, 2005; Fitzgerald et al., 2009; Waldien et al., 2018; Yonkee et al., 2019; Lease et al., 2021). The low-temperature sensitivities (<240 °C) of fission-track and (U-Th)/He dating methods are ideally suited to constraining cooling and exhumation through the middle and upper crust and the timing of brittle crustal deformation. Therefore, low-temperature thermochronometry can robustly quantify the timing, slip rates, and magnitudes of extension along detachment faults (Stockli, 2005) and can help illuminate how these structures influence the thermal nature of continental crust during large-magnitude extension (e.g., Ketcham, 1996). Although numerous structural studies have focused on the timing, rates, and magnitudes of slip along CREC detachment faults (e.g., Reynolds and Spencer, 1985; Howard and John, 1987; Davis, 1988), including studies leveraging low-temperature thermochronometry (e.g., Foster and John, 1999; Singleton et al., 2014; Prior et al., 2016), significant uncertainties continue to surround the role of crustal inheritance in the nature of CREC extensional structures, particularly their relationship to Mesozoic extensional and contractional structures and the location of the southwestern margin of the CREC. In this study, we present the first low-temperature thermochronometric data from the Riverside and Big Maria Mountains of southeastern California to constrain their Miocene structural history and exhumation along the southwestern margin of the CREC, a structurally critical region that overlaps with both the Mesozoic extensional domain of the Border rift system and the Cretaceous contractional belt of the Maria fold-and-thrust belt.

The continental crust of the western United States is underlain by the Mojave cratonic block (1.6–1.8 Ga) variably intruded by 1.4 Ga granites (Goodge and Vervoort, 2006). Throughout the late Neoproterozoic and early Paleozoic, the western edge of the North American craton developed a westward-thickening sedimentary wedge. This mechanical stratigraphy consists of early siliciclastic strata and massive, thick shelf and slope carbonates and shales that influenced the later Mesozoic and Cenozoic contractional architecture (e.g., Constenius et al., 2003; DeCelles, 2004). This Proterozoic basement and Paleozoic–Mesozoic cover sequence extended into the presentday southeast California region (e.g., Hamilton, 1982, 1987; Stone et al., 1983; Laubach et al., 1989).

Prior to Basin and Range dismemberment, a complex sequence of early Mesozoic rifting and contractional deformation affected the intra- and backarc region of the Cordilleran subduction-convergent margin system in North America (e.g., Yonkee and Weil, 2015; Fitz-Díaz et al., 2018). Jurassic arc magmatism related to the onset of the Cordilleran subduction-convergent margin from ca. 170 to 155 Ma is recorded in voluminous ca. 165 Ma granite and granodiorite plutonic rocks in the southeastern California and west-central Arizona region (e.g., Hamilton, 1982; Flansburg, 2022; Stone et al., 2022), contemporaneous with the earliest flareup of Sierra Nevada arc magmatic rocks (e.g., DeCelles et al., 2009; Yonkee and Weil, 2015). Middle Jurassic arc magmatism is also recorded as remnants of the Nazas arc in the Mexican orogen section of the Cordillera to the south (e.g., Dickinson and Lawton, 2001; Bartolini et al., 2003; Lawton and Molina-Garza, 2014; Fitz-Díaz et al., 2018). Thick siliciclastic successions interbedded with Nazas volcanic rocks suggest that arc magmatism was synchronous with opening of Middle to Late Jurassic extensional basins as part of the Border rift system along the Mexico-U.S. border, extending westward from the Gulf of Mexico into presentday southeastern California, where Late Jurassic NE-SW-directed extension in the southern Basin and Range initiated formation of the McCoy Basin (e.g., Fitz-Díaz et al., 2018). Although detrital geochronology of the first, basal deposits of the McCoy Mountains Formation suggests opening of the extensional basin in the Late Jurassic, >90% of the section consists of middle to Late Cretaceous strata that support a foreland basin depositional system associated with the Maria fold-and-thrust belt (Barth et al., 2004; Spencer et al., 2011), part of the Late Cretaceous fold-and-thrust system of the Cordilleran margin (Spencer and Reynolds, 1991).

The Cretaceous Cordilleran margin is characterized by a switch in tectonic style from the thin-skinned Sevier fold-and-thrust system to the further inboard, thick-skinned Laramide fold-and-thrust system. This change in structural style was driven by shallowing of the subducting Farallon plate, likely due to subduction of younger, hotter crust of the eastern conjugate of the Shatsky Rise (e.g., Liu et al., 2010). In between these two structural styles, the southwestern United States experienced a period of Early Cretaceous extension that has been attributed to crustal delamination (e.g., Hodges and Walker, 1992; Fletcher et al., 1995; Wells and Hoisch, 2008). The middle to Late Cretaceous was also characterized by mafic and leucogranitic magmatism, often congruent with extensional processes, in the Mojave and western margins of the presentday CREC (e.g., Miller et al., 1996; Wells and Hoisch, 2008; Economos et al., 2021).

The ~9000 km² Maria fold-and-thrust belt is defined as the Mesozoicaged thick-skinned fold-and-thrust belt that extends eastward from just south of the Iron Mountains in southeastern California to Granite Wash and the Harquahala Mountains in westcentral Arizona (Reynolds et al., 1986; Spencer and Reynolds, 1990) (Fig. 1). This E- W- trending vestige of the late Mesozoic fold-and-thrust system of western North America (e.g., Reynolds et al., 1989; Richard et al., 1994; DeCelles, 2004; Yonkee and Weil, 2015, and references therein) exposes contractional structures that involve rocks that were part of the stable North American craton and the cratonal equivalents of the Paleozoic passive-margin system until the longlived Cordilleran subduction-convergent margin initiated in the mid-Jurassic (ca. 170–160 Ma).

The Maria fold-and-thrust belt preserves Middle Jurassic, Late Cretaceous, and Paleocene Cordilleran deformation from subgreenschist facies to amphibolite facies conditions (Frost and Martin, 1982; Hamilton, 1987; Laubach et al., 1989; Spencer and Reynolds, 1990; DeCelles, 2004; DeCelles et al., 2009; Salem, 2009) that was variably involved in the Miocene large-magnitude extension of the CREC (e.g., Spencer and Reynolds, 1989, 1990, 1991). The McCoy Basin sedimentary package crops out in the McCoy Mountains (California), the Palen Mountains (California), and the Harcuvar Mountains (Arizona), recording Late Jurassic–Early Cretaceous NE-SW-directed extension. Extensional structures from the formation of the McCoy Basin likely influenced the geometries of the Late Cretaceous Maria fold-and-thrust belt. In contrast to the eastverging Late Jurassic–Early Cretaceous Sevier belt structures to the north of this region, Maria fold-and-thrust belt structures are SW vergent, although the Big Maria and Plomosa Mountains exhibit localized NE vergence (Ellis, 1982; Hamilton, 1982; Steinke, 1996, 1997; Svihla, 2003). The Paleozoic Grand Canyon stratigraphic sequence was intensively deformed and imbricated by Mesozoic basementinvolved thrusts that accommodated crustal shortening in the Maria fold-and-thrust belt (Frost and Martin, 1982). In the Moon Mountains (Arizona), Dome Rock Mountains (Arizona), and Harquahala Mountains, the Maria fold-and-thrust belt is dominated by thrust faults that carry Proterozoic and Mesozoic basement rocks in their hanging walls and imbricate moderately to highly attenuated Mesozoic and Paleozoic sedimentary sections in their footwalls (Knapp, 1989; Knapp and Heizler, 1990; Boettcher, 1996; Boettcher and Mosher, 1998). These sedimentary packages are exposed in tight to isoclinal, recumbent to overturned mapscale sheath folds in the footwall of basement-involved thrusts (e.g., Boettcher and Mosher, 1998). In the Big Maria and Little Maria Mountains (California), shortening was accommodated not by discrete brittle thrusts, but via more diffuse ductile shearing and complex folding, refolding, and extreme ductile attenuation (e.g., Ellis, 1982; Hamilton, 1987; Svihla, 2003; Salem, 2009). In contrast, the Riverside and Harquahala Mountains preserve structurally higher and less attenuated sections of the Maria fold-and-thrust belt structures that are located in the footwall of discrete Mioceneaged CREC detachment faults (Lyle, 1982; Reynolds et al., 1989; Stern, 1998). Zircon and apatite U-Pb geochronology from pegmatite dikes that crosscut the overturned Big Maria syncline in the Big Maria Mountains suggest contractional motion in the Maria fold-and-thrust belt ceased by ca. 85 Ma (Salem, 2009; Flansburg et al., 2019, 2021a; Flansburg, 2022; Flansburg and Stockli, 2022).

The thickened crust of the Cordilleran convergent margin began to collapse during the latest Cretaceous to Eocene with the development of the earliest Basin and Range metamorphic core complexes in presentday British Columbia, Idaho (USA), and western Montana (USA) (Wernicke, 1981; Coney and Harms, 1984; Spencer and Reynolds, 1989; Constenius, 1996; Colgan and Henry, 2009; Foster et al., 2010; Howlett et al., 2021), and the collapse may have been driven by the excess of stored gravitational potential energy of the hinterland (e.g., Dewey, 1988; Rey et al., 2001; DeCelles et al., 2009; Vanderhaeghe, 2012).

Previous ⁴⁰Ar/³⁹Ar, K-Ar, apatite U-Pb, and titanite U-Pb geothermochronology of the Big Maria Mountains suggests significant cooling through midcrustal temperatures (~350–550 °C) in the Late Cretaceous (ca. 80 Ma) to earliest Eocene (ca. 58 Ma) (Martin et al., 1982; Hoisch et al., 1988; Salem, 2009; Flansburg et al., 2019, 2021a, 2021b; Flansburg, 2022; Flansburg and Stockli, 2022). In the Riverside Mountains, a Late Cretaceous (ca. 70 Ma) extensional pulse is also suggested by ⁴⁰Ar/³⁹Ar cooling ages (Stern, 1998), whereas earlier ca. 105–125 Ma cooling through midcrustal temperatures is suggested by apatite U-Pb geothermochronology (Flansburg et al., 2021b; Flansburg, 2022; Flansburg and Stockli, 2022). Late Cretaceous exhumation is documented to the west of the CREC in the Iron Mountains (e.g., Wells et al., 2002; Wells and Hoisch, 2008) and to the southwest in the Mojave region (e.g., Boettcher and Walker, 1993; Wells et al., 2005; Wong and Gans, 2009; Wells et al., 2012). Notably, Cretaceous ages are proposed for the origins of some midcrustal fabrics exposed in CREC metamorphic core complexes (e.g., Anderson et al., 1988; Singleton and Wong, 2016).

Most metamorphic core complexes in the Colorado River extensional corridor (CREC) exhumed deepseated crustal rocks in the latest Oligocene through Miocene, ca. 23–14 Ma, and were accompanied by synextensional magmatism (e.g., Anderson and Rowley, 1981; Howard and John, 1987; Spencer and Reynolds, 1989, 1991; Miller et al., 1998; Singleton and Mosher, 2012; Singleton et al., 2014; Prior et al., 2016, 2018). Metamorphic core complex formation was accompanied and followed by slip along highangle normal faults (e.g., Howard and John, 1987) that formed the grabens and horsts that characterize the Basin and Range topographic expression. Dispersed Late Miocene to recent (≤10 Ma) dextral strikeslip faulting in the CREC (e.g., Singleton, 2015; Singleton et al., 2019; Mavor, 2021) is related to distributed strain from the change in the stress field along the Pacific–North American plate boundary associated with the impingement of the Mendocino fracture zone and initiation of the San Andreas fault zone (Atwater, 1970; Atwater and Stock, 1998).

The CREC developed as a series of stacked NE-directed low-angle normal faults that succeeded in unroofing ductilely deformed midcrustal rocks in their footwalls and contain SW- to W-dipping tilt blocks of Miocene volcanic and volcaniclastic rocks overlying Mesozoic and Proterozoic basement rocks in their hanging walls (Howard and John, 1987; Spencer and Reynolds, 1991; Faulds et al., 2001). The Whipple domain, named after the regional Whipple detachment fault in the Whipple Mountains metamorphic core complex, accommodated 40–60 km of extension (e.g., Spencer and Reynolds, 1989, 1991; Faulds et al., 2001). In this domain, many studies have utilized the expansive 18.8 Ma Peach Spring Tuff (Ferguson et al., 2013) as a marker horizon to date or constrain detachment faulting (e.g., Dorsey and Becker, 1995; Miller et al., 1998) or have utilized low-temperature thermochronometry to constrain detachment fault slip history (e.g., Foster and John, 1999; Singleton et al., 2014; Prior et al., 2016). The western boundary and breakaway fault of the CREC has been identified to be within the Old Woman Mountains–Iron Mountains region of southeastern California (e.g., Howard and John, 1987; Spencer and Reynolds, 1991).

Despite multiple episodes of broadly NE-SW-directed contraction and extension in the southern Basin and Range province since the mid-Mesozoic, the southwestern margin of the CREC allows the opportunity to investigate the relationship between the Mesozoic and Miocene history of this region (Fig. 1). In southeastern California, the Riverside and Big Maria Mountains expose Maria fold-and-thrust belt contractional structures in the footwalls of CREC detachment faults. Understanding the timing and nature of CREC large-magnitude extension in these ranges is crucial to reconstructing their pre-Miocene geometry and elucidating their earlier, Mesozoic to early Cenozoic tectonic evolution.

South of the well-studied Whipple Mountains (Fig. 1), the Riverside Mountains also consist of a CREC-related metamorphic core complex (Carr and Dickey, 1980), with a footwall composed dominantly of Mesoproterozoic–Paleoproterozoic basement gneisses (Flansburg, 2022) characterized by mylonitic stretching lineations consistent with top- to- the-NE shear (Fig. 2), whereas the eastern parts of the range contain late Mesozoic thrust sheets composed of Paleozoic–Mesozoic metacarbonate and siliciclastic rocks that are part of the Maria fold-and-thrust belt (Carr and Dickey, 1980; Lyle, 1982; Carr, 1991; Stern, 1998). In contrast to other CREC metamorphic core complexes, footwall fabrics are likely Mesozoic in age rather than Miocene (Flansburg et al., 2021a, 2021b; Flansburg, 2022; Flansburg and Stockli, 2022). These rocks in the footwall of the Riverside Mountains core complex are separated by the Riverside low-angle detachment fault from Cretaceous and Jurassic granitic rocks and Miocene conglomerates in the hanging wall (Stern, 1998; Flansburg, 2022). Locally, the detachment is exposed as an oxidized fault mirror above cataclasites dipping toward the SSW (Fig. 2) and forms a ~12- km- long NE-SW-trending corrugation and thus runs sinuously from the northwest to the southcentral part of the range where Big Wash dissects the higher topography (Fig. 3A). In the southwestern portion of the Riverside Mountains, the exposed detachment fault is SW dipping, suggesting that it was passively back-rotated during isostatic doming of the detachment fault, whereas in the northeastern part of the range, the detachment dips to the NE and is buried under Quaternary sediment (e.g., Stern, 1998). Previous work in the CREC region has largely ignored the Riverside Mountains, likely due to the range’s complex geology, small size, and limited road accessibility, especially when compared with the more extensively studied neighboring mountain range—the Whipple Mountains. There are no published low-temperature thermochronometric data or detailed geochronology available for the Riverside Mountains. The latest detailed mapping was by Carr and Dickey (1980) and Stern (1998), who interpreted the exposed detachment fault as an extension of the Whipple detachment fault. Mylonitic and gneissic foliations across the range dip dominantly moderately to shallowly to the NE and SW, although some foliations dip steeply to the NNE, NNW, and SSE, highlighting the domal nature of the footwall corrugation. Stretching lineation in the footwall corrugation plunges toward the SW close to the exposed detachment fault and plunges toward the NE in the northeastern portion of the range (Fig. 2C).

Located just west of the Colorado River along the center of the California-Arizona border (Fig. 1), immediately south of the Riverside Mountains, the Big Maria Mountains preserve the structural relationships of the Late Cretaceous Maria fold-and-thrust belt, dominated by the overturned SW-vergent Big Maria syncline (Hamilton, 1987; Stone et al., 2022). These Mesozoic contractional structures are crosscut by numerous postkinematic Late Cretaceous leucogranitic dikes (e.g., Salem, 2009), and the entire range has been dissected by Late Miocene to Pleistocene high-angle normal faults and strike-slip faults associated with the Eastern California shear zone (e.g., Dorsey et al., 2021; Mavor, 2021; Bennett et al., 2022; Stone et al., 2022). The Big Maria syncline notably preserves the folded metamorphic equivalents of the Paleozoic–Mesozoic Grand Canyon sedimentary section that on the inverted limb are locally attenuated by ductile shearing to <100 m thickness (Hamilton, 1987; Salem, 2009). The recognition of these prominent Mesozoic structures in the Big Maria region has made it the prime location to study Jurassic rifting associated with the McCoy Basin (e.g., Barth et al., 2004), Cretaceous contraction (e.g., Hamilton 1982, 1987; Hoisch et al., 1988; Svihla, 2003; Salem, 2009), and Late Cretaceous–Paleocene extension (Flansburg and Stockli, 2022).

Previous ⁴⁰Ar/³⁹Ar hornblende ages from Jurassic granodioritic rocks and Cretaceous pegmatite dikes suggest cooling of the Big Maria Mountains through ~450 °C from ca. 80 to 70 Ma (Martin et al., 1982; Hoisch et al., 1988; Salem, 2009), and ⁴⁰Ar/³⁹Ar white mica and biotite ages range from ca. 60 to 50 Ma, suggest cooling of the range to temperatures <300 °C by the early Eocene. However, no published low-temperature thermochronometric data exist to constrain the timing and geometric evolution of Miocene extension along the southwestern margin and breakaway region of the CREC in southeastern California. Although abundant ca. 10–7 Ma dextral strikeslip and highangle normal faults have been documented (Mavor, 2021), no evidence for Miocene detachment faulting has been observed within the Big Maria Mountains. Previous mapping studies have thus interpreted a detachment fault flanking the northern side of the range as an extension of the Riverside detachment fault (e.g., Stone et al., 2022). Other studies have placed the range entirely outside the CREC or in the upper plate and correlated the bounding breakaway fault system of the CREC between the Big Maria and Riverside Mountains (e.g., Hamilton, 1984; Spencer and Reynolds, 1991; Yin, 1991).

This study presents detailed low-temperature apatite and zircon (U-Th)/He data from the Riverside and Big Maria Mountains to constrain their progressive spatiotemporal unroofing history. Further, these data elucidate the structural architecture and location of the southwestern breakaway of the CREC as well as suggest the influence of pre-Cenozoic crustal structures on CREC formation.

The zircon and apatite (U-Th)/He thermochronometric systems are characterized by partial retention zones (PRZs), or isotherms above which He completely diffuses out of the crystal lattice and below which He is completely retained. For zircon, the (U-Th)/He PRZ is bracketed by the ~200–140 °C isotherms (e.g., Reiners et al., 2002, 2004; Wolfe and Stockli, 2010) and for the apatite (U-Th)/He PRZ, by the ~80 °C and ~40 °C isotherms (e.g., Wolf et al., 1998; Farley, 2000; Ehlers and Farley, 2003). At typical crustal geothermal gradients of ~30 °C/km, these two PRZs correspond to ~5–7 km and 2–3 km depth, respectively, and can be utilized to constrain exhumation in contractional and extensional settings (e.g., Ehlers, 2005; Stockli, 2005).

For this purpose, we collected a total of 31 samples for (U/Th)/He analyses from the Big Maria Mountains (n = 10) and the Riverside Mountains (n = 21) in rangescale transects (Figs. 3A and 4A) parallel to the regional slip direction in the CREC, ~050° ± 10° as recorded in the Whipple Mountains (Davis et al., 1980; Davis and Anderson, 1991) and ~040°–048° in the Plomosa Mountains (Strickland et al., 2018; Thacker et al., 2020). This orientation is parallel to the NE-trending footwall corrugation and coincidentally parallel to the older, mylonitic stretching lineation (Flansburg, 2022) preserved in the Riverside Mountains with a ~045° trend (Fig. 2C). Sample information and locations, along with mean apatite and zircon (U-Th)/He ages, are given in Table S1 in the Supplemental Material¹ and shown on Figures 3 and 4. In the Riverside Mountains, a ~15 km sampling transect (transect 1, Fig. 3) follows the topographic trend of the main footwall corrugation and the Miocene slip direction (~045°) (Fig. 2C). The Big Maria Mountains are characterized by a NW-SE topographic range trend due to Late Miocene strikeslip faulting, but we collected a transect (transect 2, Fig. 4) spanning the maximum NE-SW spatial extent of the range in the Miocene slip direction (~045°) parallel to transect 1 from the Riverside Mountains (Fig. 3A). Footwall samples from the Riverside Mountains consist of dominantly Mesoproterozoic–Paleoproterozoic and Jurassic quartzofeldspathic gneisses. Hanging-wall samples from the Riverside Mountains comprise mainly Jurassic and Cretaceous granitic and granodioritic gneisses. Samples from the Big Maria Mountains were collected from Jurassic granites and granodioritic gneisses as well as Mesoproterozoic–Paleoproterozoic quartzofeldspathic gneisses and quartzmica (calc-)schists in the southwestern Big Maria Mountains. We obtained apatite and zircon via standard mineral separation procedures, including crushing, grinding, water and heavy-liquid density separation, and Frantz magnetic separation at the University of Texas at Austin.

For zircon (ZHe) and apatite (AHe) (U-Th)/He analyses, 6 grains were picked from rocks with igneous protoliths and 9–12 grains were picked as singlegrain aliquots from rocks with sedimentary protoliths to capture possible intrasample variations in effective uranium (eU) concentrations among detrital grains. Individual grains were wrapped in platinum foil packets and repeatedly laser heated to 1050 °C (apatite) and 1300 °C (zircon) to ensure total degassing. Liberated ⁴He was purified and spiked (³He) in customdesigned stainless steel ultrahigh vacuum (UHV) extraction lines and analyzed using a Blazers Prisma QMS-200 quadrupole mass spectrometer. Degassed zircon were removed from the platinum foil packets and both degassed apatite and zircon were then spiked with a 7 N nitric solution with a ²³⁵U-²³⁰Th-¹⁴⁹Sm isotopically enriched tracer for isotope dilution measurements. After spiking, zircon were dissolved using a twostep pressurevessel acid digestion involving a 72-hour HF-HNO₃ pressurevessel digestion at ~220 °C and subsequently a 24-hour (180 °C) HCl acidification, following the procedure detailed in Wolfe and Stockli (2010). Samples were analyzed in 5% nitric solution by inductively coupled plasma mass spectrometry (ICP-MS) to determine parent nuclide concentrations (²³⁸U, ²³⁵U, ¹⁴⁷Sm, and ²³²Th) using a ThermoElement2 ICP-MS with a 50 μl microconcentric nebulizer.

The (U-Th)/He ages were calculated using a standard morphometric alphaparticle ejection correction (Farley et al., 1996), assuming a homogeneous parent nuclide concentration. Standard errors of 8% (zircon) and 6% (apatite) were applied to individual aliquots based on the laboratoryinternal longterm reproducibility of the Fish Canyon Tuff zircon and Durango apatite standards, respectively. Zircon with anomalously old ages (>3σ) were excluded from arithmetic mean sample age calculations and are likely attributable to partial grain loss prior to dissolution or parent nuclide zonation (e.g., Hourigan et al., 2005). Apatite with anomalously old AHe ages (i.e., older than corresponding sample ZHe age) were excluded from AHe mean sample age calculations and are shown as gray data points on Figures 3 and 4. Many of the anomalous AHe ages are characterized by relatively high eU concentrations or unusual Th/U (Fig. S1, Table S5, see ^{footnote 1}) and are likely attributable to Th- and U-rich microinclusions that were not detected during preanalysis microscope screening (180×). All ZHe raw data can be found in Table S2 and reduced ZHe data can be found in Table S3, and raw and reduced AHe data can be found in Tables S4 and S5, respectively. All (U-Th)/He analyses were collected at the UTChron Laboratory at the University of Texas at Austin.

We present 21 ZHe mean ages and 14 AHe mean ages from 17 footwall and 4 Hanging-wall samples across the Riverside detachment fault (Fig. 3). We plotted ZHe and AHe aliquot and mean ages versus distance in the detachment slip direction (NE, or ~045°) (Fig. 3B). Zircon from the Riverside Mountains yielded (U-Th)/He dates that range from ca. 75 Ma to ca. 18 Ma, decreasing from southwest to northeast (Fig. 3). Effective uranium concentrations are ~10–1300 ppm and Th/U is <0.1–1.2 (Fig. S1, Table S3, see ^{footnote 1}). In the hanging wall of the Riverside detachment fault, ZHe mean ages are consistently latest Cretaceous to Paleocene (ca. 72–58 Ma) (Fig. 3). In contrast, samples from the footwall of the Riverside detachment show a systematic age reduction toward the northeast. From southwest to northeast, we observe Paleocene to early Eocene (ca. 62–53 Ma) ZHe mean ages within ~1.5 km of the detachment in the dominant footwall corrugation (samples 19RM04 and 20RM20 on Fig. 3A), a ~2- km- wide zone of Eocene to Early Miocene (ca. 50–20 Ma) ZHe mean ages, and an expansive third region in the northeastern section of the Riverside Mountains that exhibits exclusively Miocene ZHe mean ages (ca. 21–19.5 Ma), with the exception of sample 19RM14 (29.9 ± 1.8 Ma), the northeasternmost sample (Fig. 3).

Apatite (U-Th)/He analyses yielded mean AHe ages ranging from ca. 32 to 10 Ma across the Riverside Mountains, although most sample ages cluster between ca. 25 and 20 Ma (Fig. 3). Individual aliquot ages show significant age dispersion due to inclusions or poor apatite quality and range from ca. 55 Ma to ca. 3 Ma (Fig. S2; Table S5). The youngest AHe mean age (9.7 ± 0.6 Ma) is in the northeastern corner of the range (sample 19RM13), and AHe mean ages generally decrease systematically from southwest to northeast along the footwall corrugation (Fig. 3B). Hanging-wall samples exhibit slightly older mean AHe ages of ca. 28 Ma (Fig. 3). Apatite eU ranges from 1.3 to 76 ppm (except one aliquot having eU = 141 ppm), and Th/U ranges from 0.1 to 6.6 (Fig. S1; Table S5).

Individual ZHe mean ages from nine samples from the Big Maria Mountains range from ca. 22 to 17 Ma (Figs. 4 and 5). At first glance, it seems that ZHe mean ages decrease toward the northeast (Fig. 4), with duplication observed in the two northeasternmost samples (Figs. 4 and 5) that are separated from the rest of the Big Maria Mountains transect (transect 2) by a mapped eastdipping highangle normal fault. However, all ZHe mean ages overlap within standard deviation (Fig. 4), casting doubt on any real spatial trend. Individual aliquots of zircon from the Big Maria Mountains yield a slightly larger range in ZHe age, from ca. 27 Ma to ca. 15 Ma (Fig. 4B; Fig. S2; Table S3). Effective uranium concentrations range from ~10 to 1100 ppm, and Th/U ranges from ~0.1 to 0.8 except one aliquot with Th/U = 1.4 (Fig. S1; Table S3).

The same samples analyzed for ZHe were also analyzed for AHe ages, except for sample 19BMM31, which did not contain any unbroken or inclusionfree apatite suitable for analysis. All AHe aliquots overlap in (U-Th)/He date within 2σ uncertainty, ranging from ca. 26 Ma to 11 Ma (Fig. 4B; Fig. S1) and yield mean AHe age (eight samples) of ca. 22–13 Ma (Fig. 4). Like the ZHe mean age pattern, there is no discernible spatial AHe age trend and arithmetic mean ages overlap within standard deviation (Fig. 4). Apatite eU ranges from ~3 to 47 ppm, and Th/U is ~0.2–7.6 (Fig. S1; Table S5).

Our detailed (U-Th)/He low-temperature thermochronometric data allow us to interpret the timing and magnitude of exhumation along the southwestern margin of the CREC. Understanding this Miocene extensional history provides the groundwork for reconstructing pre-CREC extensional and contractional structures and relationships preserved in the Riverside and Big Maria Mountains (Flansburg, 2022; Flansburg and Stockli, 2022). Following our discussion of (U-Th)/He data from both ranges, we suggest that the Riverside and Big Maria Mountains were exhumed between ca. 22 and 16 Ma by two structurally separate top-to-the-NE detachment faults. Extension along the western boundary of the CREC was likely controlled by pre-existing SW- and NE-verging structures of the Maria fold-and-thrust belt.

The Riverside Mountains show exhumation through the ZHe PRZ from the Late Cretaceous to Miocene. In the hanging wall of the Riverside detachment fault (Fig. 3), exclusively Late Cretaceous ZHe mean ages show that these rocks were at temperatures cooler than ~180 °C by the beginning of the Cenozoic. Footwall rocks close to the Riverside detachment (e.g., sample 19RM04, Fig. 3) preserve Paleocene ages, indicating they were also cooler than 180 °C prior to CREC-related Miocene slip on the Riverside detachment fault.

The range of Late Cretaceous (ca. 72 Ma) to Miocene (ca. 11 Ma) ZHe mean ages across the Riverside Mountains suggests the exhumation of a fossil zircon (U-Th)/He PRZ in the footwall of the Riverside detachment fault and is reflected in the northeastward-decreasing Paleocene to Early Miocene ZHe ages seen in the middle of the Riverside footwall corrugation as well as within the Cretaceous thrust sheets in the southeastern exposure of the range (Fig. 3). Because there is a systematic decrease in ZHe ages parallel to the Hanging-wall transport direction, we can track the exhumation of footwall rocks as they successively passed through the ZHe PRZ and calculate the slip rate along the Riverside detachment fault by utilizing the slope of a line regressed through ZHe age along distance in the slip direction (e.g., Wells et al., 2000; Stockli, 2005; Brichau et al., 2006). This technique requires three assumptions: (1) isotherms do not change position (Ketcham, 1996), (2) cooling ages are from samples that occupy similar depths below the detachment slip surface, and (3) displacement along the detachment fault is the sole driver of the cooling of footwall rocks. (U-Th)/He and other low-temperature thermochronometric systems help to constrain the geothermal gradient and the first assumption, but the second and third assumptions require no thinning of the hanging wall. This is a difficult assumption to satisfy, given that ZHe ages would be the result of combined horizontal (faultparallel) and vertical (isostatic) motion under the rollinghinge model (e.g., Wernicke and Axen, 1988). In this case, the slope would calculate the migration rate of the rolling hinge rather than slip along the fault; therefore, we present our determined rates as apparent slip rates out of caution. Using the robust linear regression in Isoplot (Ludwig, 2009), an apparent slip rate with 95% confidence intervals of 3.2 +1.0/−1.1 km/m.y. (Fig. 5) can be calculated from ZHe ages structurally below the ZHe PRZ in the Riverside Mountains. This slip rate is consistent with rates of ~2–7 km/m.y. in the Harquahala, Harcuvar, Chemehuevi, and Buckskin-Rawhide detachment faults (Foster et al., 1993; John and Foster, 1993; Singleton et al., 2014; Prior et al., 2016).

In general, ZHe ages from the footwall of the Riverside detachment systematically decrease toward the northeast in the direction of extensional slip and parallel to the detachment corrugation (Fig. 3). Previous structural studies show that the Riverside detachment fault is domed and backrotated to the southwest (Carr and Dickey, 1980; Stern, 1998). Therefore, the exposure observed in the Riverside Mountains (Fig. 2) is not the most recently active part of the detachment’s slip surface, which is now under Quaternary sedimentary deposits to the north and northeast of the range. To account for the doming of the Riverside detachment fault and the general northeast-younging trend observed in ZHe mean ages, there are two possibilities. The rolling-hinge model of metamorphic core complex formation (Axen et al., 1995), in which slip is progressively younger in the slip direction, could account for the northeastward decrease in ZHe age as well as provide a mechanism for isostatic uplift of the lower plate of the Riverside detachment fault. However, there is a shift in AHe mean ages from ca. 22–19 Ma to ca. 16–15 Ma in the northeastern end of the footwall corrugation. Rather than a rolling hinge, a secondary breakaway that soled into the lower (now domed) strand of the Riverside detachment fault in the central part of the NE-oriented footwall corrugation could account for this AHe mean age shift (Fig. 6). This secondary breakaway could also explain the presence of Miocene AHe mean ages in the hanging wall of the Riverside detachment as well as the younger AHe mean ages in the northeastern section of the footwall corrugation, ca. 16.5–14.7 Ma (Fig. 3), which overlap with AHe mean ages in the Big Maria Mountains (Fig. 4). Removal of the secondary breakaway’s hanging wall would have allowed for unloading and isostatic rise of the Riverside detachment fault and the SW-dipping, back-rotated surface observed across the range today (Figs. 3 and 6) and would account for the overlap in ZHe and AHe mean ages in the northeastern end of the range (Fig. 3B), which is suggestive of a secondary, structurally higher fault that rapidly unroofed these rocks through both the ZHe and AHe PRZs at approximately the same time. This geometry is like that observed in the Whipple Mountains metamorphic core complex to the north, where a secondary breakaway at ca. 14.5 Ma is constrained by basalts deposited onto an already exposed footwall (Dorsey and Becker, 1995). Other possibilities that could account for the lack of systematic, or directionally related, AHe age variation in both the Riverside and Big Maria Mountains are (1) repetition of cooling ages across incised strands of the detachment into the footwall, as inferred by Stockli et al. (2006) in the Whipple Mountains; (2) hydrothermal fluid–assisted resetting of cooling ages; and (3) brittle extensional attenuation of the hanging wall. We favor erosion-related, isostatically driven uplift of the Riverside Mountains to account for the prevalence of Middle Miocene ages there, especially given that there is no direct evidence of post-detachment highangle normal faults, detachment incisement strands, or pervasive latestage hydrothermal alteration that would result in uniform AHe mean ages across the footwall of the range.

All of the ZHe mean ages from the entire span of the Big Maria Mountains are ca. 20 Ma, with all mean ages overlapping within 2σ uncertainty and standard deviation. If the Riverside detachment fault was responsible for the exhumation of the Big Maria Mountains, we would expect to see a similar Cretaceous to Miocene age pattern as in the Riverside Mountains, or if the Big Maria Mountains resided entirely in the hanging wall of the Riverside detachment, we would expect pre-Miocene ZHe ages across the range. These data rather suggest that the entirety of the Big Maria Mountains was exhumed through the ZHe PRZ coevally between ca. 22 and 17 Ma; therefore, we propose that the range was exhumed from underneath the Riverside Mountains by a structurally lower detachment and observe two northeast-younging trends in ZHe mean age that are separated by mapped faults (Fig. 6). We can calculate an apparent slip rate for this detachment, which we call the Big Maria detachment, using the same methods and assumptions as utilized for the Riverside detachment (detailed in the Timing and Magnitude of Exhumation of the Riverside Mountains section). Apparent slip rates for the Big Maria detachment are 3.0 +1.7/−1.8 km/m.y. and 2.7 +0.0/−6.1 km/m.y. with 95% confidence intervals (Fig. 5). These apparent slip rates were calculated from two northeast-younging ZHe age domains across the Big Maria Mountains, rather than one single domain, because the two northeasternmost samples of the transect are separated from those to the southwest by mapped highangle normal faults: the SW-dipping Quién Sabe fault (Richard, 1993; Dembosky and Anderson, 2005), an unnamed NW-striking normal fault, and the SE-dipping Slaughter Tree Wash fault (Mavor, 2021) (Fig. 4A). Further, the repetition in (U-Th)/He ages may suggest that the northeastern end of the transect either is offset by a buried detachment mapped just north of the range but offset by the Slaughter Tree Wash fault on the eastern end (Hamilton, 1984) or represents an excised portion of the Big Maria detachment similar to what is suggested for the Whipple detachment to the NNE (e.g., Stockli et al., 2006). It is also possible that the buried detachment mapped by Hamilton (1984) is the Big Maria detachment proposed in this work, but one important distinction with our work is that the Hamilton (1984) interpretation, also present in the Stone et al. (2022) map, correlates this buried detachment with the Riverside detachment to the north (Fig. 4A), whereas we favor two distinct, stacked detachments.

Though we present apparent slip rates for the Big Maria detachment, it is important to note that all ZHe ages across the Big Maria Mountains overlap within standard deviation as well as 2σ uncertainty. This means that if we calculate an apparent slip rate across the entire range, slip rates approach infinity, suggestive of extremely rapid or unilateral uplift across the range (Fig. 5B). Further, AHe ages overlap with each other and with the ZHe ages, also implying fast slip rates. Together, this suggests that these data violate the assumptions required to calculate an apparent slip rate. By interpreting stacked detachment faults in this region (Fig. 6), we inherently violate two of the assumptions that underlie these slip rate calculations: (1) that there is no vertical advection of the footwall, and (2) that there is no upperplate extension. This implies that these slip rates are extremely fast (i.e., approaching infinity) and that these slip rates should be used with caution; we use these rates to compare apparent slip on the southwestern margin of the CREC to slip along detachments toward the northeast.

Another scenario for exhumation of the Big Maria Mountains would be exhumation of the range from underneath a steeply NE-dipping normal fault, inducing southwestward backrotation of a tilt block in the footwall. However, this hypothesis would require decreasing (U-Th)/He ages toward the southwest, away from the highangle normal fault, rather than toward the northeast, unless it invoked largemagnitude southwestward backrotation to account for nearsimultaneous exhumation through the ZHe and AHe PRZs and thus generating overlapping AHe and ZHe ages. Coupled with the ~15 km lateral spatial extent of our data across the range, giving the backtilted block a long wavelength, the overlap in AHe and ZHe ages makes this scenario unlikely.

The fossil ZHe PRZ and relict Cretaceous ZHe cooling ages preserved in the Riverside Mountains are consistent with the Riverside Mountains being located structurally higher than the Big Maria Mountains to the south. The detachment responsible for exhumation of the Big Maria Mountains would be located just to the NNE of the range, buried by the presentday dune field and the Quaternary sediments of Big Wash (Fig. 4A; “Big Maria detachment” on Fig. 6), which drains the central and southern portions of the Riverside Mountains. This stacked detachment geometry is like that of the main detachment faults of the Colorado River extensional corridor, such as the Plomosa detachment beneath the Buckskin-Rawhide detachment in Arizona (e.g., Singleton et al., 2014; Spencer et al., 2018).

It is likely that the Riverside detachment fault is structurally lower than the Whipple detachment fault to its north, like how the Big Maria detachment is structurally lower than the Riverside detachment in this region. Because the Buckskin-Rawhide detachment is likely correlative with the Whipple detachment (e.g., Howard and John, 1987; Spencer and Reynolds, 1991), it may be that the Plomosa and Riverside detachments are also correlative within the broader CREC extensional system and structure. The relationship between the Riverside and Plomosa detachments is similar to the correlation between the Whipple and the Buckskin-Rawhide detachments, which requires a WNW jump in the CREC detachment system.

The new, lower detachment fault recognized in this study (“Big Maria detachment” in Figs. 6 and 7) would therefore be the structurally lowest detachment fault of the CREC and integral to the exhumation of rocks preserving the Cretaceous–Paleocene deformation of the Maria fold-and-thrust belt. The data we present here are the first conclusive thermochronometric evidence that the CREC breakaway exists between the Big Maria and Riverside Mountains, as first correlated by Spencer and Reynolds (1991) and Yin (1991) (Fig. 7). Accounting for ca. 10–7 Ma dextral and dextraloblique strikeslip motion in the Blythe region (Fig. 1) associated with the Eastern California shear zone and Pacific–North American plate boundary (Dorsey et al., 2021; Mavor, 2021; Bennett et al., 2022) would restore the Big Maria Mountains next to the Moon and northern Dome Rock Mountains across the Colorado River in Arizona, suggesting that the detachment exposed in the northeastern Moon Mountains, previously interpreted as the structurally lowest CREC detachment, is likely the Big Maria detachment as well (e.g., Spencer and Reynolds, 1991; Spencer et al., 2018).

The lowtemperature thermochronometry presented here suggests that the Big Maria and Riverside Mountains were exhumed via two coeval, separate detachments that initiated in the latest Oligocene to Early Miocene (Fig. 6). Rapid fault slip is documented by exhumation through the ZHe PRZ from 23.3 ± 1.8 Ma to 19.5 ± 1.2 Ma in the Riverside Mountains (Figs. 3 and 5) and from 22.1 ± 1.8 Ma to 17.0 ± 1.4 Ma in the Big Maria Mountains (Figs. 4 and 5). The Big Maria detachment is the structurally lowest detachment in the domain and forms the southwestern boundary of the CREC. The domed and backrotated expression of the Riverside detachment fault, coupled with consistent ca. 20 Ma AHe ages, is indicative of a secondary breakaway in the northeastern region of the range, consistent with the northeastward migration of strain observed in other CREC metamorphic core complexes, such as the Whipple and Buckskin-Rawhide Mountains (e.g., Dorsey and Becker, 1995; Singleton et al., 2014; Prior et al., 2018). Interestingly, the data presented here suggest that the initiation of slip along the southwestern boundary of the CREC at ca. 21–17 Ma overlaps in time and was likely coeval with ca. 22–19 Ma slip initiation along the structurally higher and larger detachment systems that formed the Chemehuevi-Whipple-Buckskin-Rawhide-Harcuvar-Harquahala, Plomosa, and Sacramento metamorphic core complexes to the northeast and north, suggesting that CREC extension was temporally synchronous over distances >250 km (Foster and Spencer, 1992; Foster et al., 1993; Scott et al., 1998; Foster and John, 1999; Carter et al., 2004, 2006; Stockli et al., 2006; Sanguinito, 2013; Singleton et al., 2014; Gans and Gentry, 2016; Prior et al., 2016).

Restoring the ZHe PRZ exposed in the Riverside Mountains to a ZHe PRZ in the Big Maria Mountains, at a minimum distance above the range and since eroded away, would approximate a minimum of ~20 km of horizontal displacement along the Big Maria detachment (Fig. 6), although there are no clear lithologic or structural offsets to further constrain this estimate and further detailed structural work is needed. Because the ZHe PRZ is not exposed in the Big Maria Mountains, a more conservative estimate of horizontal displacement would be the NE-SW transect distance across the range, ~17 km, which resides entirely within the footwall and has Miocene ZHe and AHe cooling ages. While the Big Maria and Riverside detachment faults were active at approximately the same time, this estimate is approximately half of that for displacement along the structurally higher detachments in the north, northeast, and east, where ~40–50 km displacements are recorded in the Whipple, Buckskin-Rawhide, Harcuvar, and Harquahala Mountains (e.g., Reynolds and Spencer, 1985; Faulds et al., 2001; Singleton et al., 2014; Prior et al., 2016), suggesting that these structurally lower detachments that bound the southwestern extent of the CREC accommodated less extension than meta morphic core complex–forming detachments to the northeast.

Due to the prevalence of Mesozoic and Early Cretaceous SW- and NE-verging structures in the Big Maria–Riverside region, we can begin to speculate on the importance of previous crustal fabrics on the initiation and location of lowangle normal faults that accommodate largemagnitude extension in the continental crust. The dominance of SW-NE-oriented structures and crustal fabrics traces its origins back to the Jurassic, when the McCoy Basin developed as an extension of the Border rift system. In the Big Maria and Riverside Mountains, Late Cretaceous contractional structures, including thrust faults and folds, associated with the Maria fold-and-thrust belt are both SW and NE verging (e.g., Hamilton, 1987; Stern, 1998; Svihla, 2003; Salem, 2009), and the Big Maria Mountains also preserve Late Cretaceous–Paleocene mylonitic shear zones with top-to-the-NE extensional kinematics (Flansburg, 2022; Flansburg and Stockli, 2022). Further, due to their location with respect to the Riverside detachment, our reconstruction presented in Figure 6 restores the Paleozoic metasedimentary rocks of the Riverside Mountains, buried under Cretaceous thrust faults, along the Riverside detachment prior to Miocene slip. Much of this metasedimentary succession is calcite rich (e.g., Carr and Dickey, 1980; Hamilton, 1987; Stern, 1998) and may have been an important factor in localizing strain during Miocene extension, such as interpreted in the Catalina-Rincon and Buckskin-Rawhide metamorphic core complexes (Singleton et al., 2018; Spencer et al., 2019). This long history of SW-NE-directed deformation was likely a major influence on the location and structural style of the detachment faults responsible for exhuming the Riverside and Big Maria Mountains and determining the nature and location of the western boundary of the CREC.

The (U-Th)/He thermochronology presented in this study represents the first lowtemperature thermochronometric data from the Riverside and Big Maria Mountains of southeastern California and helps constrain the exhumation of ranges to the southwest of the major Miocene detachment faults of the Colorado River extensional corridor (CREC). The Riverside Mountains preserve a fossil Miocene zircon (U-Th)/He (ZHe) partial retention zone, with footwall rocks yielding ca. 50 Ma (Eocene) to ca. 11 Ma (Miocene) ZHe mean ages that young toward the northeast, whereas the entirety of the Big Maria Mountains yields ZHe mean ages of ca. 20 Ma. Significantly, these data suggest that the Big Maria Mountains were exhumed from underneath the Riverside Mountains via a structurally lower detachment fault, the Big Maria detachment, now covered by the modernday dune field and Big Wash. ZHe ages yield apparent slip rates of ~3–4 km/m.y. for the Riverside detachment and the Big Maria detachment faults, and the Big Maria detachment accommodated a minimum of ~17 km of horizontal displacement between the two mountain ranges. Apatite (U-Th)/He age results suggest a secondary breakaway in the Riverside Mountains that allowed for isostatic uplift and backrotation of the Riverside detachment at ca. 16–14 Ma. These results provide conclusive evidence of the CREC breakaway between the Big Maria and Riverside Mountains as the southwesternmost and structurally lowest detachment fault bounding the CREC.

We wish to thank Rudra Chatterjee and Des Patterson for assistance with (U-Th)/He analyses at the UTChron Laboratory at the University of Texas at Austin, as well as Maya Ortiz and Catherine Ross for field assistance. This work benefited from conversations with Michael Wells, Richard Ketcham, Sharon Mosher, Nikki Seymour, Cullen Kortyna, and Michael Prior and thoughtful reviews from John Singleton and Stephen Reynolds. Funding for this work was provided by the Geological Society of America (graduate student grant awarded to Flansburg), the Jackson School of Geosciences at the University of Texas at Austin, and UTChron Laboratory funds (Principal Investigator: Stockli).