Structural analyses combined with U-Pb zircon petrochronology show the influence of arc magmatism on the evolution of two transpressional shear zones in the deep root of the Late Cretaceous Southern California batholith. The mid-crustal Black Belt and lower-crustal Cucamonga shear zones (eastern San Gabriel Mountains, California, USA) formed at ca. 84 Ma, shortly after a large mass of tonalite and granodiorite intruded the lower crust. Both shear zones were active until at least ca. 74 Ma and probably until 72–70 Ma. In the mid-crustal shear zone, rheological contrasts between mingling magmas localized deformation at dike margins. The deformation began as hypersolidus flow in partially crystallized dikes and then transitioned to deformation below the solidus when alternations between viscous creep and brittle faulting produced interlayered pseudotachylyte, cataclasite, and mylonite. As the dikes solidified, strain hardening drove shear zone growth and created thin (10–30 m) high-strain zones and faults that are widely spaced across ~1 km. In contrast, the lower-crustal Cucamonga shear zone was magma-starved, lacks the variety of shear zone fabrics exhibited by its mid-crustal counterpart, and formed by the reactivation of a preexisting fabric that records pure reverse displacements at 124–93 Ma. The two shear zones created a partitioned style of intra-arc transpression where sinistral-reverse (mostly arc-parallel with some arc-oblique) displacements were accommodated on moderately dipping faults and shear zones and arc-normal shortening was accommodated by coeval folds. This study shows how a magmatic surge influenced the architecture and style of Late Cretaceous transpression in the Southern California batholith, including the evolution of high-strain zones that record alternating episodes of brittle, ductile, and hypersolidus deformation. The results illustrate how magmatism localizes strain on deep-crustal faults during orogenesis and oblique convergence.
]]>The study of Cordilleran-style magmatic arcs is essential to our understanding of orogenesis and the evolution of continental lithosphere. These regions, which develop above ocean-continent subduction zones, are the factories that create new continental crust (Tatsumi, 2005;Ducea et al., 2015;Cashman et al., 2017;Collins et al., 2020). They also are primary sites where mass and heat move through the lithosphere by a wide variety of mechanisms, including compaction and porous flow (Rabinowicz and Vigneresse, 2004;Holness, 2018;Gleeson et al., 2023), brittle faulting and fracturing (Huber et al., 2011;Webber et al., 2015;Contreras-Reyes et al., 2021), buoyancy-driven flow in channels and diapirs (Whitney et al., 2004;Kruckenberg et al., 2013;Klepeis et al., 2016), and deformation in melt-enhanced shear zones (Kelemen and Dick, 1995;de Saint Blanquat et al., 1998;Webber et al., 2015;Klepeis et al., 2022). Understanding exactly how deformation and magmatism interact within these systems is important because these interactions build the arc and often link to other dynamic processes operating within convergent margins.
One of the most challenging aspects of reconstructing arc histories is determining the evolution of their deep roots. The lower crust of arcs is widely thought to be a crucible where the influx of magma and heat from the mantle combine with intra-crustal MASH (melting, assimilation, storage, homogenization) processes, leading to magma diversification and crustal recycling (Hildreth and Moorbath, 1988;Cashman et al., 2017;Ratschbacher et al., 2018;Collins et al., 2020;Hanson et al., 2022). However, these processes, and those that move melts into and out of lower-crustal MASH zones, are poorly understood. One reason for this is that much of what we know has been inferred from xenoliths (Ducea and Saleeby, 1996;Ducea, 2002;Chin et al. 2014), geophysical surveys (Huang et al., 2013;Contreras-Reyes et al., 2021;Jiang et al., 2023), and lab experiments (Rapp and Watson, 1995;Müntener et al., 2001;Getsinger et al., 2009;Ulmer et al., 2018). Whereas these studies are crucial for identifying the processes and structures that contribute to arc construction, we still lack direct information on how magmatism, deformation, and MASH processes interact with one another in different settings.
Exhumed batholiths that expose large tracts of arc crust from depths greater than ∼25 km are relatively rare. Some notable examples occur in the North Cascades (Miller et al., 2009), the southern Sierra Nevada batholith (Klein et al., 2021), Argentina's Valle Fértil (Walker et al., 2015), the Salinian Block in central California (Kidder et al., 2003), and New Zealand's Median batholith (Allibone et al., 2009;Schwartz et al., 2017Schwartz et al., , 2021;;Klepeis et al., 2019). One particularly understudied example, which forms the focus of this paper, is the Cucamonga terrane in the southeastern San Gabriel Mountains of southern California, USA (Figs. 1 and 2; May and Walker, 1989;Barth and May, 1992;Schwartz et al., 2023Schwartz et al., , 2025b)). This terrane forms part of an ∼500-km-wide segment of the Mesozoic California arc that lies between the southern Sierra Nevada batholith and northern Peninsular Ranges batholith (Fig. 1). Recent work by Schwartz et al. (2023Schwartz et al. ( , 2025aSchwartz et al. ( , 2025b) ) has shown that a Late Cretaceous surge in arc magmatism within this batholith was linked to the formation of a network of transpressional shear zones. One of these, the Black Belt shear zone (Fig. 2), localized inside a large mass of tonalite and granodiorite that intruded middle-and lower-crustal gneiss. These exposures provide an important opportunity to examine magma-deformation interactions during a magmatic flare-up at the root of the California arc.
Previous work in the southeastern San Gabriel Mountains (May and Walker, 1989;Barth and May, 1992;Schwartz et al., 2023Schwartz et al., , 2025b) also has shown that zones of mylonitic fabrics are interlayered with gneiss that records high-grade metamorphism and deep-crustal melting. Schwartz et al. (2025b) reported cataclastic fabrics that are interlayered with mylonite and pseudotachylyte, all of which were reworked at amphibolite-facies conditions. Pseudotachylyte often is interpreted to represent a solidified frictional melt produced by seismic slip on a fault (Sibson, 1975;Toy et al., 2011;White, 2012;Hawemann et al., 2019;Orlandini and Mahan, 2020;Michalchuk et al., 2023). Its presence below the classically defined seismogenic zone (>25 km depth) is thought to reflect transient, high-velocity slip in rock that was, at least momentarily, strong enough to support anomalously high differential stresses (Mahan et al., 2008;White, 2012;Regan et al., 2014;Orlandini and Mahan, 2020;Menegon et al., 2021;Michalchuk et al., 2023) or, in rarer cases, the result of ductile instabilities that locally produce melting (e.g., White, 2012). The occurrence of pseudotachylyte in the Cucamonga Block, and especially its interlayering with mylonite, raises the possibility that ductile instabilities led to localized disequilibrium melting along rheological contrasts within crystallizing plutons deep within the batholith. We explore this possibility in this study.
Here, we combine structural data with U-Pb zircon geochronology to show how the midcrustal Black Belt and lower-crustal Cucamonga shear zones formed within the root of the Southern California batholith. Our goals included: (1) to determine the internal structure, kinematics, and timing of deformation in the two shear zones; (2) to identify the processes that controlled their evolution, including the potential localizing effects of episodic magmatism and inherited structures; and (3) to determine the possible origin and significance of the different types of fabrics that make up the shear zones, including interlayered pseudotachylyte, cataclasite, and mylonite, and igneous fabrics that record hypersolidus flow. These topics are discussed in the context of new maps and cross sections (Figs. 3 and 4) that show the structure of the region and document a four-stage sequence of events (D 1 -D 4 ) that influenced its evolution.
The Southern California batholith (SCB) lies between the southern end of the Sierra Nevada batholith where it meets the Garlock fault and the northern end of the Peninsular Ranges batholith (Fig. 1). Together these three batholiths form the Mesozoic California arc. In the SCB, late Paleozoic to Mesozoic magmatism occurred in three pulses at 260-210 Ma, 160-140 Ma, and 90-70 Ma (Schwartz et al., 2023). The terminal phase peaked at 85-75 Ma (Economos et al., 2021;Schwartz et al., 2023Schwartz et al., , 2025b) ) slightly after a ca. 98 Ma pulse in the Sierra Nevada batholith and a ca. 93 Ma pulse in the Peninsula Ranges batholith (Ducea et al., 2015;Paterson and Ducea, 2015;Paterson et al., 2017). Late Cenozoic strike-slip faulting (Powell, 1993;Dickinson, 1996) has dismembered and dispersed fragments of the SCB across the Transverse Ranges, the Salinia Block (Fig. 1), and the western Mojave province.
In contrast to the Sierra Nevada batholith where pre-Mesozoic host rock is preserved in small metasedimentary pendants, large blocks of Proterozoic basement have been reported from the SCB (Barth et al., 1997(Barth et al., , 2017;;Needy et al., 2009;Nourse et al., 2020;Economos et al., 2021;Schwartz et al., 2023). In the southeastern San Gabriel Mountains (Fig. 2), the Cucamonga terrane is composed of mafic and felsic gneiss, amphibolite, quartzofeldspathic gneiss, minor marble, and paragneiss (Hsu, 1955;May and Walker, 1989;Barth and May, 1992;Barth et al., 1992). To date, no Proterozoic gneiss has been found in this unit. Instead, Schwartz et al. (2025b) found that the protolith to these rocks appears similar in age and composition to Sur Series metasedimentary rocks in the Salinia Block and may have been deposited in a late Paleozoic to early Mesozoic forearc or intra-arc basin. During the Cretaceous, they were affected by multiple phases of amphibolite-and granulite-facies metamorphism (Hsu, 1955), involving pressures of 7-9 kbar and temperatures of 700-835 °C (Barth and May, 1992;Schwartz et al., 2025b). These data show that the Cucamonga terrane formed part of the hot, lower crustal root of the SCB during the Cretaceous and, thus, represents one of only a few areas where rocks situated at depths of 25-30 km can be observed directly (Anderson, 1988;Barth, 1990;Barth and May, 1992;Needy et al., 2009).
North of the Cucamonga terrane, the San Antonio terrane is composed of a series of tonalites, diorites, and granodiorites (Fig. 2). These intrusive rocks, which originally were defined as the Tonalite of San Sevaine Lookout by Mor-ton et al. (2001), intruded the Cucamonga terrane and formed part of a suite of felsic-intermediate rocks that was emplaced into the middle crust (∼6 kbar) of the arc at 90-73 Ma (Barth, 1990;Miller and Barton, 1990;Foster et al., 1992;Miller et al., 1996;Barth et al., 2008;Chapman et al., 2017Chapman et al., , 2021;;Schwartz et al., 2023Schwartz et al., , 2025aSchwartz et al., , 2025b;;this study). The exposures of high-pressure granulites at the southern end of the Cucamonga terrane and mid-crustal plutons in the central and northern parts of the San Gabriel Mountains suggests that the two units form part of tilted crustal section that shallows to the north.
The northern and southern contacts of the Cucamonga terrane coincide with two upper amphibolite-facies shear zones called the Black Belt and Cucamonga shear zones, respectively (Fig. 2; Alf, 1948;Schwartz et al., 2025b). May and Walker (1989) concluded that these features constitute a west-directed ductile thrust system that formed during arc magmatism. We show here that this deformation formed part of a sinistral transpressional regime within the batholith. Schwartz et al. (2023Schwartz et al. ( , 2025b) ) showed that the deformation occurred during emplacement of the Tonalite of San Sevaine Lookout at
89-85 Ma and ended before cross-cutting dikes were emplaced, nominally at ca. 74 Ma.
Several other shear zones of comparable style and a similar Late Cretaceous age occur south of the Black Belt and Cucamonga shear zones. In the northern Peninsular Ranges batholith, the Eastern Peninsular Ranges mylonite zone (EPRM; Fig. 1) records west-directed thrusting during Late Cretaceous time (Simpson, 1984;Morton et al., 2014). The late Cenozoic San Jacinto fault system has broken this shear zone into segments that include, from north to south, the Santa Rosa Mylonite, the Coyote Mountain Mylonite, and the Borrego Springs shear zone. These zones record deformation and metamorphism in the upper crust under epidoteamphibolite (450-600 °C; Simpson, 1984Simpson, , 1985;;Goodwin and Renne, 1991), upper amphibolite (∼700 °C; Theodore, 1970), and upper greenschist (∼400 °C) facies conditions (Simpson, 1984), respectively. Results from titanite thermochronology and 40 Ar-39 Ar dating suggest that the Santa Rosa Mylonite formed at 90-77 Ma (Goodwin and Renne, 1991;Jiang et al., 2015).
The Western Transverse Ranges expose transpressional shear zones that formed at midcrustal depths within the batholith. The Tumamait shear zone (Fig. 1) in the Pine Mountain Block records intra-arc sinistral-reverse shearing from 77 Ma to 70 Ma that was synchronous with the emplacement of peraluminous granite sheets (Schwartz et al., 2025a). Deformation at upper amphibolite-facies temperatures of ∼700 °C was followed by folding and arc-normal thrusting on the Sawmill thrust fault at 67-66 Ma. In the central Transverse ranges, the mid-crustal Alamo Mountain and Piru Creek shear zones record sinistral strike-slip motion at 76-72 Ma, involving temperatures of 500-600 °C and pressures of 4 kbars (Bixler et al., 2025). Other Late Cretaceous shear zones include the Nacimiento fault (located west of the Tumamait shear zone, Fig. 1), which records sinistral displacements that placed Late Cretaceous granitoids of the Salinia Block against Franciscan Complex rocks of the Nacimiento Block at 75-60 Ma (Mattinson, 1990;Hall, 1991;Kidder at el., 2003;Dickinson,1983;Dickinson et al., 2005;Singleton and Cloos, 2013;Johnston et al., 2019). The Mill Canyon Window (Fig. 1) in the west-central San Gabriel Mountains also records syn-plutonic, top-to-the-NW (present-day orientation) thrusting at 73-62 Ma (Barth et al., 2019). Together, these structures form a semi-continuous belt of reverse, sinistral strike-slip, and sinistral-reverse displacements within the Late Cretaceous batholith (May, 1989;Schwartz et al., 2025a;Bixler et al., 2025).
The dominantly sinistral component of transpressional deformation in the SCB contrasts with the primarily dextral transpressional deformation that occurred in segments of the magmatic arcs in Idaho, northern Nevada, and eastern California (Tikoff et al., 2023, and references therein). Most of these latter structures occur north of the present-day location of the B A Figure 3. Maps of (A) Deer and (B) Cucamonga canyons showing the geometry of D 2 to D 4 structures in the San Antonio (pink) and Cucamonga (green) terranes. Note the wide spacing of high-strain zones within the Black Belt shear zone, some of which exhibit pseudotachylyte (p)-bearing brittle faults interlayered with mylonite, cataclasite, and ultramylonite. See text for discussion. Locations of sites sampled for geochronology are shown with numbers in white boxes. Numbers correspond to last digit(s) of sample numbers described in the text. White numbers in black boxes refer to other sites discussed in the text. Red lines with triangles are faults that form part of the late Cenozoic Cucamonga fault zone. DCF is the Demens Canyon fault. Yellow shading represents undifferentiated Quaternary sediments located south the Cucamonga fault zone. White areas are canyon drainages where most of the data were taken.
Garlock fault and include the Mojave-Snow Lake fault, the Sierra Crest shear zone, and the western Idaho shear zone, among others (Fig. 1; Fossen and Tikoff, 1993;Giorgis et al., 2017;Krueger and Yoshinobu, 2018;Trevino et al., 2021;Tikoff et al., 2023, Schwartz et al., 2025a). Dextral transpression has been interpreted to be related to the collision between the Insular superterrane and the western margin of North America at 100-85 Ma (Tikoff et al., 2023).
The San Gabriel Mountains are bounded on the north by the San Andreas fault, which has accumulated ∼300 km of dextral strike-slip displacement since ca. 5 Ma (Powell and Weldon, 1992). The southern boundary coincides with the Cucamonga thrust fault, which forms part of a network of transpressional faults that occur within a restraining bend in the San Andreas fault. This zone of thrusting is composed of at least three strands that become progressively younger to the south (Matti et al., 1982) and places metamorphic rocks of the Cucamonga terrane over alluvial fan sediments (Fig. 2). Previous work suggests that the fault zone dips 25°-43° to the N (33° on average) near the surface (Morton and Matti, 1987;McPhillips and Scharer, 2018) and steepens to 50° at depth (Cramer and Harrington, 1987). Estimates of total displacements on the strands are limited by a lack of piercing points.
A network of secondary faults within the San Gabriel Mountains has helped drive uplift of the range since the late Miocene (Spotila et al., 2002;Mere and McPhillips, 2024). The ages of most of these faults are reported by the U.S. Geological Survey and California Geological Survey (USGS and CGS, 2024) and Morton and Miller (2006). Many appear to have been active since ca. 4 Ma. The San Gabriel fault, which deforms the central part of the massif, records ∼60 km of dextral displacement since ca. 10 Ma (Crowell, 1982). These and other faults contribute to the total strain budget across the range (Mere and McPhillips, 2024). The Cucamonga thrust became the principal dip-slip structure in the system at 1.2-0.5 Ma (Matti and Morton, 1993;Blythe et al., 2000).
Faults that predate the onset of transpression include early-middle Miocene extensional and transtensional faults located within and outside the eastern San Gabriel Mountains (Nourse, 2002). On the basis of correlations with dikes and volcanic rocks, Nourse (2002) inferred a late Miocene (18-13 Ma) age for the faults and linked them to a zone of transtension that formed in a right stepover between the dextral Clemens Well and Offshore Borderland faults. In this paper, we report grabens and sinistral-normal faults that form part of this transtensional system. Wong et al. (2023) also identified a period of Miocene (21-20 Ma) extension that affected much of southeastern California and west-central Arizona. Other studies show that much of the orogenic hinterland of the western cordillera experienced extension from Eocene to Miocene time (Yonkee and Weil, 2015, and references therein).
The various late Cenozoic strike-slip faults in southern California divide the Transverse Ranges into tectonic blocks, many of which have moved relative to one another (Terres and Luyendyk, 1985;Carter et al., 1987;Luyendyk, 1991;Nicholson et al., 1994;Dickinson, 1996;Onderdonk, 2005Onderdonk, , 2007;;Meade and Hager, 2005;Marshak, 2016;Ingersoll and Coffey, 2017). Combinations of paleomagnetic data, structural data, and geodetic measurements have been used to estimate these rotations. Luyendyk (1991) reported two episodes of verticalaxis rotations for the San Gabriel Block, which lies between the San Andreas and San Gabriel faults. The first involved 53° of clockwise rotation after 21 Ma and the second 16° of counterclockwise rotation after 11 Ma (a net clockwise rotation of 37°). Both estimates were based on paleomagnetic data obtained by Terres and Luyendyk (1985; see also Hornafius et al., 1986). Because our field site lies south and east of the San Gabriel fault, it is unclear if these rotations apply to Cucamonga and Deer canyons.
Marshak (2016) used paleomagnetic measurements from middle Miocene basaltic dikes from the southeastern San Gabriel Mountains, including from our study area, to suggest that the region has rotated ∼20° counterclockwise since ca. 5 Ma. This result matches the prediction of counterclockwise rotation required to bring the San Gabriel Block to its present position from where it was near the modern-day Salton Sea prior to ca. 5 Ma (Ehlig 1981, Dillon andEhlig 1993). To our knowledge, no information has been published that suggests major tilting of the study area around a horizontal or inclined axis. We explore this possibility further in the Discussion section.
Standard field mapping techniques were used to record all observable phases of brittle and ductile deformation. These included measuring the orientations of foliations, folds, and mineral stretching lineations. Shear sense was determined using common asymmetric structures, including mica and hornblende fish, C-S fabrics, C′ shear bands, and asymmetric recrystallized tails on porphyroclasts. Crosscutting relationships among rock fabrics (S 1 , L 2 /S 2 , L 3 /S 3 ; e.g., Figs. 5A-5C and 5N), variably deformed intrusions (Figs. 5C-5J), and a set of late folds (F 4 ; Figs. 5C and 5P) were used to develop a relative sequence of deformation, metamorphism, and magma emplacement. These groupings provided a basis for comparing structures between canyons and defining the relative ages of four tec-tonic events (D 1 -D 4 ). We then used U-Pb zircon geochronology to determine the absolute age of these events.
Previous studies have shown that shear zone geometries may be complex and require 3-D kinematic treatments (Czeck and Hudleston, 2003;Jones et al., 2004;Iacopini et al., 2007;Kuiper et al., 2011;Giorgis et al., 2017;Díaz-Azpiroz, et al., 2019). The kinematic vorticity vector (W k ), for example, must be found independently of the orientations of foliations and lineations. We followed accepted methods for finding W k in the field (Díaz-Azpiroz et al., 2019), including testing for triclinic symmetry and determining displacement directions by evaluating shear indicators on multiple planes. We also looked at how fold and foliation orientations change across strain gradients in order to assess whether deformations were consistent with regional contraction (or transpression) and crustal thickening or regional extension (or transtension) and crustal thinning. This approach is possible because different flow types produce distinctive fold geometries and rotation patterns (Fossen et al., 2019). The use of strain gradients also helped us link observations made at the scale of individual outcrops to kilometer-scale structural patterns. We used observations of late Cenozoic faults to restore present-day orientations to their Late Cretaceous configurations, as discussed in the Cenozoic Block Rotations subsection in the Discussion. For areas that record mostly brittle deformation, we distinguished normal, reverse, and oblique-slip faults by measuring fault planes and slickenlines, and determined shear sense using offset layering, Riedel shears, stepped microfractures, microfaults, and shear bands.
To identify deformation that occurred in dikes and plutons prior to their full crystallization, we used criteria described by Paterson et al. (1989Paterson et al. ( , 1998)), Vernon (2000Vernon ( , 2004)), Pawley and Collins (2002), Yoshinobu et al. (2009), Turnbull et al. (2010), Webber et al. (2015), and Klepeis et al. (2016). These studies describe structures that commonly form during transitions from hypersolidus flow in melt-rich magmas to deformation near the solidus where crystal-plastic deformation occurs in the presence of migrating melts. For anatectic migmatites, we used criteria described by Hopgood (1999), Gervais et al. (2004), Vanderhaeghe (2009), Kruckenberg et al. (2011), Sawyer (2014), and Klepeis et al. (2016). These studies distinguish layered migmatites where foliation planes that pre-date partial melting are preserved (metatexites) from
examples where foliation planes have been destroyed and replaced by syn-anatectic flow structures (diatexites). We corroborated fieldbased interpretations of melt-present deformation using U-Pb zircon geochronology and used the results to define periods where deformation was synchronous with magmatism and/or crustal anatexis.
We conducted U-Pb analyses on 11 zirconbearing samples (Table 1) to determine the timing of magmatism, metamorphism, and deformation in the Black Belt and Cucamonga shear zones. Some zircons exhibited interior core and rim domains that reflect a history of deformation and metamorphism at amphibolite-to granulite-facies conditions. Where possible, we used zircon trace-element concentrations and ratios, particularly U (ppm), U/Th, Gd/Yb, and Dy/Yb to distinguish metamorphic age populations (see Supplemental Material 1 for more detail).
To determine the ages of different stages of deformation, we sampled variably deformed intrusions and their hosts and divided them into pre-, syn-, and post-kinematic categories with respect to rock foliations and lineations. The pre-kinematic variety is represented by dikes and plutons that were emplaced prior to deformation (e.g., Figs. 5G and 5H) and thus provide a lower (older) limit on deformation events. Syn-kinematic intrusions (e.g., Fig. 5C), where dikes both cut foliation planes and are tightly folded within them, were especially useful for dating deformation that was synchronous with magmatism. Undeformed, post-kinematic dikes (e.g., Fig. 5I) were emplaced after deformation had ceased, thus placing an upper (younger) age limit on events. This approach, and comparisons with published ages, allowed us to bracket the ages of rock fabrics.
The analytical methods used in this study closely follow those outlined in Kylander-Clark et al. (2013) and are described in detail by Schwartz et al. (2025b). U-Pb ratios were collected at the University of California, Santa Barbara, USA, using a Nu Plasma multi-collectorinductively coupled plasma mass spectrometer (MC-ICPMS) with a Photon Machines 193 ArF excimer laser with HelEx cell. Spot size and frequency were 35 µm and 4 Hz, respectively. The primary standard (91500) was accessed every 10 analyses to correct for in-run fractionation of Pb/U and Pb isotopes. A secondary standard (Temora-2) was analyzed every ∼10 analyses to assess data reproducibility. Uncertainties are reported as 2SE internal calculated from Iolite and IsoplotR (Paton et al., 2010;Vermeesch, 2018). We assigned a 2% uncertainty to all dates for interlab comparisons to account for the reproducibility of standards.
Zircon ages are reported using the 206 Pb/ 238 U date for analyses <1100 Ma, and the 207 Pb/ 206 Pb date for those >1100 Ma. For the former, discordance is calculated as the percent difference 1 Supplemental Material. Materials include U-Pb isotope data for zircons from the Cucamonga and San Antonio terrains and cathodoluminescence images of representative zircons from sites 51b and 20 in the Cucamonga granulite. Please visit https:// doi .org /10 .1130 /GSAB .S.28840865 to access the supplemental material; contact editing@geosociety .org with any questions.
Magma influence on fault and shear zone evolution Geological Society of America Bulletin, v. 136, no. XX/XX 9 Cathodoluminescence images were obtained using an FEI Quanta scanning electron microscope before and after ablation to evaluate analyzed areas and compare them with growth textures (e.g., Fig. S1). Where possible, we targeted all growth domains and report 207 Pb/ 206 Pbcorrected 206 Pb/ 238 U ages of texturally homogeneous populations. In a few cases where a laser spot overlapped multiple domains, we report the data in tables but did not consider them in weighted mean calculations. Concordia plots and error-weighted average ages are summarized in Tables 1 and S1.
Trace elements (Table S1) were measured simultaneously with U-Pb isotopes by LA-sector field (SF)-ICPMS using Zr as the internal standard and nominal values of 43.14% Zr. Trace element data were reduced using Iolite (Paton et al., 2010(Paton et al., , 2011) ) and concentrations calculated relative to U.S. National Institute of Standards and Technology Standard Reference Materials (NIST)-612 as a primary standard. BHVO-2G was analyzed as a secondary standard to assess reproducibility of the data. We determined precision by repeated analysis of basaltic glass NIST-612. Long-term precision ranges from 2.8% to 17.1% and is <7.5% for all analyzed elements excluding P and Tm. Accuracy, as measured for basaltic glass BHVO-2G, was better than 5% for all elements except P and Hf. All laser-ablation data including standard information are contained in the Supplemental Material. For zircon, model Ti-in-zircon temperatures were calculated using the Ferry and Watson (2007) calibration. Samples of granulite-facies rocks from the Cucamonga terrane contain rutile, allowing us to estimate the activity of TiO 2 to be one. For samples that lack rutile, we assume a value of 0.6 based on the presence of ilmenite.
At Cucamonga and Deer canyons, the Cucamonga terrane is composed of granulite-facies paragneiss, migmatite, felsic and mafic orthogneiss, and marble. The paragneiss includes interlayered quartzite, biotite schist, calcareous gneiss, amphibolite, and quartzofeldspathic gneiss. Mafic gneiss includes garnet-bearing amphibolites and other rocks rich in biotite, garnet, orthopyroxene, clinopyroxene, and plagioclase. Felsic granulites, including charnockites (orthopyroxene-bearing gneissic granites) and deformed tonalites, are interlayered with the mafic gneisses. All of these rocks were intruded by moderately deformed to undeformed tonalite, granodiorite, and granite dikes.
Most of the Cucamonga terrane displays a penetrative gneissic foliation (S 2 ) whose composition varies depending on the protolith. In paragneiss, S 2 is defined by aligned quartz, plagioclase, biotite, potassium feldspar, and garnet. In mafic rocks, it is composed mainly of plagioclase, hornblende, orthopyroxene, clinopyroxene, and biotite ± garnet. In felsic gneiss, it tends to be mylonitic, where both quartz and plagioclase display evidence of grain size reduction. A penetrative down-dip mineral stretching lineation (L 2 ) plunges gently to moderately to the NW on NW-dipping S 2 planes (Figs. 3A, 3B, and 6A) throughout the section. In mafic layers and paragneiss, L 2 is defined by aligned hornblende and plagioclase; in felsic layers, it is defined by aggregates of plagioclase, potassium feldspar, and quartz. In some areas, especially in paragneiss, the L 2 /S 2 fabric forms lenses that envelop an older set of granulite-facies gneissic foliations (collectively referred to as S 1 ) that lack a mineral lineation (Fig. 5A). Inside the lenses, S 1 and veins of granitic leucosome that parallel it are tightly folded (F 2 ) and truncated by S 2 . These veins and other areas of patchy leucosome provide textural evidence of high-grade metamorphism and crustal anatexis (M 1 ) that occurred prior to the formation of the dominant L 2 /S 2 fabric. This early history (D 1 ), and the events that produced the L 2 /S 2 fabric (D 2 ), occur exclusively within the Cucamonga terrane and are absent in the San Antonio terrane.
Sense-of-shear indicators within L 2 /S 2 are present throughout the section. In felsic layers, they include asymmetric hornblende fish and asymmetric quartz tails on plagioclase grains. In paragneiss, the lozenges where S 2 envelops S 1 also are asymmetric (Fig. 5A). S 2 at the margins of these lenses forms mylonitic shear bands. Everywhere they were observed, the plane of maximum asymmetry shown by these indicators occurred on surfaces oriented approximately parallel to the down-dip L 2 mineral lineation and perpendicular to S 2 , indicating approximately monoclinic strain symmetry. The indicators all show a consistent top-up-to-the-SE (present geometry) reverse sense of shear parallel to L 2 . This pattern defines a 2.5-km-thick zone of distributed lower-crustal shortening that accommodated reverse-style displacements (D 2 ).
Where paragneiss and mafic orthogneiss in the Cucamonga terrane are rich in biotite, they tend to be migmatitic. These migmatites provide evidence that granulite-facies metamorphism and partial melting (M 2 ) of the lower crust accompanied D 2 shortening. At site 25, for example, a metatexite preserves undeformed to weakly deformed accumulations of leucosome that cut across S 2 foliation planes in host rock and fill dilatational sites, including tension gashes and interboudin partitions (Fig. 5B). The leucosome shows higher abundances of potassium feldspar, plagioclase, and quartz, and lower abundances of biotite and other mafic minerals compared to the matrix. These accumulations display a variety of shapes and sizes, ranging from diffuse globular patches to dike networks with irregular, scalloped margins, to veins with straight, planar boundaries. Because they both cut across foliation and are caught up in the deformation, the leucosome textures indicate that partial melting accompanied D 2 deformation. Together, the D 1 and D 2 events record a history of high-T metamorphism and crustal anatexis that occurred both prior to (M 1 ) and during (M 2 ) a major lower-crustal thickening event (D 2 ).
Regional architecture. The Cucamonga shear zone is exposed at the southern ends of Cucamonga and Deer canyons, where it is superimposed on older D 2 structures in the Cucamonga terrane (Fig. 3). Its northern boundary is marked by the appearance of mylonite and tight folds (F 3 ) of L 2 /S 2 and is gradational, reflecting a southward increase in the intensity of D 3 deformation. The southern boundary coincides with the Cucamonga thrust fault. Both boundaries dip moderately to the NW (Fig. 4). Shear zone widths are variable, ranging from ∼500 m in Deer Canyon to >1.5 km in Cucamonga Canyon (Fig. 3). This variability most likely reflects late Cenozoic faulting.
Inside the shear zone, zones of mylonite are widely spaced (Fig. 3). The intermediate-strain zones between them are characterized by penetrative gneissic and schistose S 3 foliations that deform S 2 /L 2 and display biotite and quartz mineral lineations (L 3 ) on their surfaces. These L 3 lineations plunge gently to the ENE and WSW (Fig. 6B). Their variability reflects late F 4 folds that deform the L 3 /S 3 fabric (Figs. 4 and 5C).
North of the Cucamonga shear zone and within gneissic rock of the Cucamonga terrane, the effect of D 3 deformation is visible but relatively weak. The rocks in this zone display a pen-etrative S 2 foliation that locally was reactivated by D 3 but typically without producing mylonite or tight folds (F 3 ). This reactivation is indicated by the presence of an oblique L 3 biotite mineral lineation that is superimposed on a down-dip L 2 hornblende lineation on S 2 surfaces (Fig. 7A). The two lineations are easily distinguished on the basis of orientation and composition (compare L 2 and L 3 lineations in Figs. 6A and 6B; Fig. 7A). These observations show that D 3 deformation affected the entire lower-crustal section, although the effect of D 3 between the Black Belt and Cucamonga shear zones is weak.
Superimposed on all L 2 /S 2 and L 3 /S 3 structures in the Cucamonga terrane are two sets of folds (F 4 ). These folds show two orthogonal orientations that create a regional dome-andbasin interference pattern. The dominant set is composed of tight, asymmetric folds (Fig. 5C) that plunge to the NNE and SSW (Fig. 6B) and verge to the SSE (Fig. 4). The subordinate set is upright, plunges to the north and south (Fig. 6B), and forms open arches. Both sets lack an axial planar foliation, and doubly plunging, disharmonic folds are common. This complexity most likely reflects the inherited structure of the Cucamonga terrane, which is structurally and compositionally heterogeneous.
Deformation and magmatism. Tonalite, granodiorite, and granite dikes intruded the Cucamonga shear zone, although in volumes that were much less than in the San Antonio terrane. Sites 20 and 23 (Fig. 3A) are important localities because they preserve evidence that dikes were emplaced synchronously with F 4 folding. At site 23, a thick felsic dike cuts across S 3 foliations and intrudes the F 4 folds (Fig. 5C). Synchronous dike emplacement and folding is indicated by the dikes both cutting the F 4 folds and being deformed by them. Elsewhere, including at site 51b (Fig. 3B), undeformed (post-kinematic) veins and dikes cross cut S 3 (Fig. 5D). These relationships indicate that minor magmatism occurred both during formation of the shear zone and after it, although the Cucamonga shear zone lacks the variety of magmatic and subsolidus fabrics that form an integral part of the Black Belt shear zone.
Kinematics. A kinematic analysis of D 3 deformation in the Cucamonga shear zone was facilitated by an abundance of sense of shear indicators, including asymmetric quartz and biotite pressure shadows on mafic and felsic clasts, C′ shear bands, asymmetric mica fish, and minor shear zones. We estimated the orientation of the kinematic vorticity vector (W k ) and displacement direction for each high-strain zone by evaluating the degree of asymmetry on multiple surfaces regardless of the orientation of L 3 and S 3 . The results show an approximately monoclinic strain symmetry with a D 3 kinematic vorticity vector that lies within S 3 and is oriented perpendicular to L 3 (Fig. 7A). The majority of shear sense indicators record a top-to-the-SW, sinistralreverse sense of displacement, except where F 4 folds cause L 3 to plunge to the SW (Fig. 6B). The kinematic significance of the F 4 domes-andbasins, together with the F 4 folds that deform the Black Belt shear zone, is discussed in the following section under the heading "Kinematics."
Regional architecture. Structurally on top (i.e., north) of the Cucamonga terrane is a series of variably deformed tonalites, diorites, and granodiorites of the San Antonio terrane (Figs. 3 and 4). This unit is sheeted where individual dikes and plutons were emplaced approximately parallel to one another, forming
A B Figure 11 shows data restored to their Late Cretaceous orientations.
a regional igneous layering that dips moderately to the NW (Fig. 4). At its southern end, an intrusive relationship with the older Cucamonga terrane is preserved, along with a 200-400-m-thick contact aureole characterized by migmatite and hornfels textures (Fig. 3). Unlike the older migmatites that either predated or accompanied formation of L 2 -S 2 , granitic leucosome in the younger migmatites that help define the contact aureole cross cut L 2 -S 2 . The Black Belt shear zone is superimposed on this intrusive boundary, penetrating ∼1 km northward into the San Antonio terrane (Figs. 3 and 4). The presence of the contact aureole, along with the intrusive relationship, suggests that the total amount of displacement accommodated by the shear zone was relatively small (a kilometer or two).
Inside the Black Belt shear zone, a series of thin (10-30-m-thick), widely spaced high-and intermediate-strain zones are separated from one another by thick low-strain zones. The lowstrain zones preserve igneous layering and linear (L > S) magmatic flow fabrics (Figs. 3, 4, 5E, and 5F). High-strain zones (e.g., Figs. 5H and 5J) are defined by the presence of mylonitic and ultramylonitic foliations (S 3 ) with quartz, biotite, and plagioclase mineral stretching linea-tions (L 3 ) on their surfaces. The mineral lineations plunge obliquely to the ENE and WSW on S 3 planes that, on average, dip moderately to the NW (Fig. 6C). The doubly plunging lineations reflect late folds (F 4 ) that form a regional dome-and-basin pattern in both the San Antonio and Cucamonga terranes (Figs. 3, 6B, and 6C). Intermediate-strain zones (e.g., Fig. 5G) display protomylonitic foliations (S 3 ) with mineral stretching lineations (L 3 ) that are similar in composition and orientation as those in high-strain zones. These intermediate zones also preserve igneous layering and lack the pronounced grain size reduction observed at higher strains (e.g., Fig. 5G). The general parallelism between igneous layering and S 3 in all three zones, regardless of strain intensity, suggests that the wide spacing of the mylonites reflects an inherited igneous architecture created by the sheeted dikes.
At the southern end of the shear zone, where it deforms rocks of the Cucamonga terrane (site 13, Fig. 3A), a L 3 /S 3 is superimposed on L 2 /S 2 in a narrow belt only a few tens of meters thick. At this locality, L 3 /S 3 locally is mylonitic and envelops lenses of tightly folded L 2 /S 2 structures that are truncated by S 3 (Fig. 5N). The L 3 /S 3 fabric is easily distinguished from L 2 /S 2 on the basis of orientation, mineral assemblage, kinematics, and the presence of mylonite and crosscutting relationships. The narrow width of the zone compared to the total thickness of the shear zone shows that D 3 deformation is preferentially localized within the San Antonio terrane.
Deformation and magmatism. Intrusive relationships and magmatic flow fabrics are prevalent both inside and outside the Black Belt shear zone. Site 30 (Fig. 3A), for example, shows elliptical enclaves of tonalite in a coarse, dioritic matrix (Fig. 5E). The enclaves are composed of plagioclase, hornblende, and biotite that show grain sizes an order of magnitude smaller than those in the diorite, with the exception of a few large crystals. Matrix plagioclase forms thin (≤1 cm) accumulations that surround and penetrate the enclaves by different degrees. A few megacrysts are frozen at enclave boundaries, recording an exchange of material prior to the complete crystallization of the two bodies. Fine-grained granodioritic dikes with internal flow foliations cut across both units (Fig. 5E). The dike margins are undulatory, suggesting that they intruded an incompletely crystallized host. Patches of biotite-rich granitic leucosome cut across the dikes and exchanged crystals with
the dioritic host. This latter relationship suggests that granitic melt was able to pool within a crystal-rich mush after emplacement of the granodiorite dikes. All of these textures show that the rocks that make up the San Antonio terrane were characterized by different effective viscosities as they intruded and mingled prior to their full crystallization. These differences created rheological contrasts between magma batches that most likely resulted from variations in grain size, composition, temperature, and melt concentration (cf. Blake and Fink, 2000;Webber et al., 2015). Site 4 (Fig. 3A) shows that hypersolidus deformation localized along the margins of mingling magmas. This low-strain site exposes a dark gray tonalite intrusion with scalloped margins that shares a magmatic flow fabric (L 3 /S 3 ) with a light-colored granodiorite intrusion (Fig. 5F). The fabric in the granodiorite is composed of tiled, tabular plagioclase phenocrysts and aligned biotite that parallels a coarsegrained hornblende and biotite fabric in the adjacent tonalite (Fig. 5F). A few plagioclase and quartz aggregates in the granodiorite also are elliptical and record evidence of grain size reduction by crystal-plastic recrystallization. Coarse, unfoliated granitic apophyses emanating from the granodiorite intruded the tonalite, forming veins that truncate the L 3 /S 3 fabric (Fig. 5F). The presence of the veins shows that, at the time they formed, the tonalite was crystalrich and capable of fracturing. Away from the contact, L 3 /S 3 merges with less deformed areas. These textures illustrate that deformation localized along the margins of comagmatic dikes as a result of a rheological contrast between the two units.
Site 9 (Fig. 3A) shows an example of L 3 /S 3 in zones of intermediate-strain where granodioritic dikes and their tonalitic hosts had mostly crystallized. At this site, the deformation resulted in tight folds (F 3 ) of igneous layering (Fig. 5G) and produced a penetrative S 3 foliation that cuts across both units and parallels the axial planes of the folds. The foliation is protomylonitic where grain size reduction was achieved by crystalplastic recrystallization. However, igneous textures, represented by coarse, tabular plagioclase also are preserved. A few meters away, in a high-strain zone (Fig. 5H), these relic igneous textures are destroyed in granodioritic dikes that have been completely recrystallized into mylonite. Despite these transitions in the degree of recrystallization, the orientations of mineral lineations (L 3 ) and foliations (S 3 ) remain parallel in low-, high-, and intermediate-strain domains (Fig. 4A). The pattern suggests that D 3 deformation initiated during hypersolidus flow in mingling tonalite and granodiorite bodies and continued to develop as the rocks transitioned below the solidus where crystal-plastic processes produced protomylonitic and mylonitic foliations.
Undeformed, post-kinematic dikes are common. Site 1 (Fig. 3A) shows an example of an undeformed granodioritic dike that intruded a deformed tonalitic host (Fig. 5I). The dike displays margins that are mostly straight with a few thin apophyses and a few undulatory segments, suggesting that, at the time it intruded, the host was a viscous, crystal-rich mush. The presence of these dikes, along with the highly deformed ones, shows that magma emplacement continued throughout the life of the shear zone. The continuous or episodic emplacement of the dikes, combined with a heterogeneous distribution of high-strain zones, explains the variety of magmatic and subsolidus deformation textures preserved in the shear zone.
Interlayered fault rocks. In addition to magmatic textures, high-strain zones in the Black Belt shear zone preserve evidence of both brittle and ductile styles of deformation. Mylonite is interlayered with thin (≤1 m), coplanar zones of cataclasite and pseudotachylyte (Figs. 5J and 5K). In areas of the highest strain, ultramylonite and discrete faults with mineral striae (biotite and quartz streaks) also are present. All of these layers parallel the margins of dikes and magmatic layering in tonalite, diorite, and granodiorite hosts (Figs. 5J and 5K).
Textural relationships among the different brittle and ductile layers preserve evidence that cataclasis and mylonitization alternated in time as the Black Belt shear zone formed. The mylonites are strongly foliated (S 3 ) and lineated (L 3 ) and display biotite and hornblende shear bands and recrystallized, asymmetric quartz tails on rotated plagioclase and potassium feldspar porphyroclasts. Where cataclastic layers escaped most of the later mylonitization, the cataclasites contain angular, fractured shards of mainly plagioclase and potassium feldspar within a fine-grained biotite-rich matrix (e.g., Fig. 5L). The matrix in these cataclasites displays a weak foliation (S 3 ) defined by aligned biotite and hornblende that is similar in orientation to the mylonitic foliation (S 3 ). A weak biotite mineral lineation (L 3 ) also is present. Away from mylonite, the cataclasites lack the asymmetric, recrystallized quartz tails and shear bands that are characteristic of mylonitic layers and tend to display clasts that are more equant and less aligned than in the mylonite (e.g., Fig. 5M). Pseudotachylytes are also strongly foliated with biotite microlites and dynamically recrystallized quartz layers defining the foliation. Survivor clasts are elliptical to rounded in shape and are dominated by plagioclase. Survivor clasts are commonly wrapped by the foliation, preserving a sense of shear in the pseudotachylyte seams that is kinematically compatible with the sense of shear in the surrounding mylonitic fabric. Figure 5L shows an example where an thick seam of pseudotachylyte interlayered with cataclasite broke into fragments that were incorporated into an adjacent mylonite, forming asymmetric clasts. The asymmetry records an oblique, top-up-to-the-SW (present-day geometry), sinistral-reverse sense of motion that matches the shear senses recorded by both the brittle faults and the mylonitic portions of the shear zone (kinematics described in the following section). Cretaceous post-kinematic dikes that cut these fabrics, and the F 4 folds that deform them, provide additional evidence that these layers formed as part of the Cretaceous Black Belt shear zone rather than from an episode of late Cenozoic thrust faulting. The lower boundary between the cataclasite and the mylonite is narrow and undulatory (Fig. 5L) whereas the upper boundary is gradational (Fig. 5M) and merges into a zone of ductile flow where mylonitic foliations (S 3 ) envelop lenses of older cataclasite. These textures suggest that, at this site, ductile flow leading to mylonitization occurred after cataclasis and melting.
Elsewhere, the opposite relationship occurs where cataclastic textures are superimposed on older mylonite. In these latter areas, elliptical clasts of plagioclase with asymmetric tails of quartz and plagioclase are broken into angular clasts within a fine-grained, recrystallized matrix (Fig. 5K). These mutual overprinting relationships suggest a type of non-steady progressive deformation where strain rates and other flow parameters changed over time. The interlayering and kinematic compatibility among the different fabrics indicate that transient ductile rupture events leading to the formation of pseudotachylyte occurred between episodes of viscous creep (Miranda et al., 2023a).
Kinematics. The preservation of well-defined L 3 mineral stretching lineations and an abundance of sense-of-shear indicators allowed us to determine the kinematics of flow in the Black Belt shear zone and to evaluate whether it formed by regional extension (or transtension) or regional shortening (or transpression). To estimate the orientation of the kinematic vorticity vector (W k ) and displacement direction, we evaluated the degree of asymmetry on multiple surfaces for each highand intermediate-strain zone, regardless of the orientation of L 3 and S 3 . Everywhere, the plane showing the most asymmetry occurred on surfaces oriented parallel to L 3 and perpendicular to S 3 (Fig. 7B). This pattern persists regardless of rock type and includes areas (e.g., site 13) where D 3 reactivated S 2 foliation planes in the Cucamonga terrane (Figs. 5N and 5O). At site 13, the reactivation produced an oblique L 3 biotite mineral lineation that is superimposed on a downdip L 2 hornblende lineation (Fig. 7A).
This result establishes an approximately monoclinic strain symmetry for the shear zone with a D 3 kinematic vorticity vector that lies within S 3 and is perpendicular to L 3 (Fig. 7B). Furthermore, the majority of high-and intermediate-strain zones record a consistent top-to-the-SW, sinistral-reverse sense of displacement. In a few places in Cucamonga Canyon, gentle F 4 folds reorient L 3 so that it plunges to the SW (Fig. 6C), yielding an apparent top-down-tothe-SW (apparent sinistral-normal) sense of displacement.
The F 4 folds that deform all D 2 and D 3 structures provide another means of evaluating the kinematic significance of deformation. These folds show two orthogonal orientations that create a regional dome-and-basin interference pattern. The dominant set is composed of tight, asymmetric folds (Fig. 5P) that plunge to the NNE and SSW (Fig. 6C) and verge to the SSE (Fig. 4). The subordinate set is upright, plunges to the north and south (Fig. 6B) and forms open arches. Both sets lack an axial planar foliation. Their effect on the D 3 shear zones is illustrated by the 3-D perspective diagram shown in Figure 8. The diagram shows the curvature of the L 3 lineations and the pattern of orthogonal fold sets. It also illustrates that a component of vertical thickening accompanied horizontal shortening, resulting in the domes and basins. This geometry allows us to rule out regional extension and transtension, which would result in different fold geometries (Fossen et al., 2019).
Another check on the kinematics is the sense of slip recorded on brittle thrust faults that display seams of pseudotachylyte interlayered with mylonite and cataclasite (Fig. 3). These faults occur in the highest strain sites (Sites 2, 7, 45, and 47; Fig. 3) and display biotite and quartz mineral streaks that plunge obliquely to the NE on gently moderately dipping fault planes (Fig. 7C). Offset layering and stepped, crescentshaped fractures show consistent top-up-to-the-SW, sinistral-reverse displacements that match the sense of shear recorded in the adjacent mylonite (Fig. 7D). They, along with L 3 /S 3 , are cut by Cretaceous dikes, providing further evidence that the high-strain zones record sinistralreverse displacements.
A series of grabens and sinistral transtensional faults deform rocks of both the San Antonio and Cucamonga terranes. These are preserved best at the northern end of Deer Canyon and in thick marble layers in Cucamonga Canyon (Fig. 9). Two orientations are evident. The first formed grabens that strike to the NNE and includes the Demens Canyon fault. The second formed sinistral transtensional faults that strike WNW and includes the Stoddard Canyon fault. Both sets are similar to other NNE-striking sinistral and oblique-normal faults located outside the San Gabriel Mountains, including the ancestral San Antonio Canyon fault, Raymond Hill fault, and Santa Monica-Malibu Coast fault. On the basis of correlations with other Miocene normal and transtensional faults (Nourse, 2002;Wong et al., 2023) we assign an early Miocene age to these grabens.
Superimposed on the grabens is a group of E-and NE-striking reverse and oblique-reverse faults that parallel the Cucamonga thrust fault (Fig. 9). These reverse faults record small displacements (a few tens of meters) where they deform intrusive contacts. Some reactivate older normal faults, others reactivate gneissic S 2 foliations. Close to the Cucamonga thrust they form a small fold-thrust belt that imbricates gneisses within the Cucamonga shear zone (Fig. 9). These minor faults represent precursors of the Cucamonga thrust and record Pliocene-Quaternary activity (USGS and CGS, 2024;Morton and Miller, 2006). Their occurrence suggests that secondary faults in the hanging wall of the Cucamonga thrust fault helped accommodate range-front uplift.
New U-Pb zircon ages and geochemistry (Fig. 10; Tables 1 and S1) helped us resolve the absolute ages of superposed magmatic, metamorphic, and deformation events in the SCB. The results establish that transpression was synchronous with magmatism and high-grade metamorphism in the Cucamonga and Black Belt shear zones.
Three samples from site 21CM51b (Fig. 3B) were especially useful for determining the age and duration of deformation (D 3 ) in the Cucamonga shear zone at Cucamonga Canyon. This site exposes a garnet granulite orthogneiss (sample 51b-Host) that is cut by two sets of leucocratic veins (Fig. 5D). The older set (sample 21CM51b-V1) is migmatitic, shares a penetrative S 3 foliation with its host, and contains metamorphic garnet. Quartz in the vein is recrystallized and elongated parallel to S 3 . A younger vein set (sample 21CM51b-V3) is undeformed, garnet-absent, and cuts both V1 and S 3 .
Sample 21CM51b-Host yielded nine xenocrystic, oscillatory-zoned zircons (Fig. S1A) with a weighted average age of 136.25 ± 2.02 Ma (mean square of weighted deviates [MSWD] = 1.9) and average temperature conditions of ∼772 °C (Figs. 10A and 10B). There is Pb loss apparent in some grains. Metamorphic rims on two zircons gave a mean age of 88.76 ± 2.73 Ma (MSWD = 1.9) and an average temperature of ∼777 °C (Fig. 10B). Xenocrystic zircon cores (n = 6) from sample 21CM51b-V1 produced a mean age of 128.33 ± 1.36 Ma (MSWD = 6.5) and average temperatures of 780 °C, which are similar to ages obtained from the host gneiss (Figs. 10C and 10D). Luminescent rims (n = 18; Fig. S1B) from this latter sample yielded a mean age of 86.25 ± 0.45 Ma (MSWD = 1.8) and an average temperature of 749 °C, which are similar to the rim ages from the host gneiss (Fig. 10D). Both samples show a flattening of HREE/MREE (heavy and medium rare earth element) patterns indicative of garnet growing in equilibrium with zircon crystallization. The results place a lower limit of ca. 86 Ma on the age of D 3 in the Cucamonga shear zone at Cucamonga Canyon. They also suggest that an episode of high-temperature, garnet granulitefacies metamorphism (M 2 ) occurred at 91-86 Ma as emplacement of the Tonalites of San Sevaine Lookout and D 3 deformation in the shear zone both were beginning. This result is in agreement with data presented by Schwartz et al. (2025b) who showed that zircons from the Cucamonga terrain record partial melting at garnet granulite conditions during this period (their fig. 10).
The results from sample 21CM51b-V3 provided an upper limit on the age of D 3 in the shear zone. The sample was collected from an undeformed, garnet-absent leucocratic vein that cuts across S 3 foliation planes (Fig. 5D). Xenocrystic cores (Fig. S1C) show a mean age of 143.95 ± 2.38 Ma (MSWD = 4.6; n = 2) and an average temperature of 768 °C (Figs. 10E, 10F). Metamorphic rims (n = 5) show a mean age of 77.42 ± 0.73 Ma (MSWD = 2.5) and an average temperature of ∼694 °C (Fig. 10F). This age is similar to other undeformed granodioritic dikes in the Cucamonga and San Antonio terranes. The results indicate that M 2 garnet-granulite metamorphism and D 3 at this site ended prior to ca. 77 Ma and was followed by cooling and lower temperature metamorphism. All three samples from site 51, thus, constrain D 3 deformation in the Cucamonga shear zone at Cucamonga Canyon to have occurred during the 86-77 Ma interval.
Sample 061022-20 helped us to establish the timing of F 4 folding in the Cucamonga shear zone. The site is composed of mylonitized tonalite, paragneiss, and amphibolite that display a penetrative S 3 foliation deformed by disharmonic folds (Fig. 5C). Sample 061022-20 is from a suite of syn-kinematic (with respect to F 4 ) felsic dikes that were emplaced into the axial plane of the F 4 folds (Fig. 5C). The dike both cuts the folds and is deformed by them. The sample contained xenocrystic zircons (Fig. S1D), the oldest of which yielded an average age of 260 ± 4.14 Ma (MSWD = 2.2) and average temperatures of ∼725 °C (Figs. 10G and 10H). The youngest population yielded a weighted average age (n = 21) of 77.17 ± 0.40 Ma (MSWD = 2.7) and average temperatures of 719 °C (Fig. 10H). This result suggests that F 4 folding occurred through ca. 77 Ma and overlapped temporally with D 3 deformation. The age matches those of other 77-74 Ma dikes that were emplaced within both the San Antonio and Cucamonga terranes.
Samples 22CM59 and 22CM60 were collected from granodiorite and tonalite dikes, respectively, that cut an S 2 foliation in banded paragneiss of the Cucamonga terrane outside the Cucamonga shear zone (Fig. 3B). The former sample is undeformed and yielded zircons with a weighted mean average age of 76.8 ± 0.3 Ma (n = 17; MSWD = 6.6; Figs. 10I and 10J) and average temperatures of ∼750 °C. This age is similar to those of other undeformed granodioritic dikes across the study area, including within the Black Belt and Cucamonga shear zones, and suggest that they all form part of a coeval igneous suite. Sample 22CM60 (Fig. 3B) is from a mylonitic tonalite dike that intruded parallel to S 2 in the banded paragneiss. The tonalite dike yielded a weighted mean average age of 78.0 ± 0.2 Ma (MSWD = 4.2) for low U zircons (n = 21; Figs. 10K and 10L). High U zircon rims (n = 8) from this same sample yielded a weighted mean average age of 75.1 ± 0.4 Ma (n = 8; MSWD = 2.3; Fig. 10L) and average temperatures of ∼729 °C. This latter age matches those from granodioritic dikes, including the undeformed ones. The gneissic foliation likely formed as D 3 deformation localized in the crystallizing dike as the nearby Cucamonga shear zone continued to deform.
Site 3 in Deer Canyon (Fig. 3A) exposes a coarse-grained, foliated (S 3 ) tonalite host (sample 060722-3A) intruded by undeformed granodiorite dikes (sample 060722-3B). The foliation in the host is protomylonitic, contains garnet, and occurs in a relatively low-strain site. The dikes are fine-grained, biotite-bearing, and cut across a penetrative S 3 foliation. Sample 060722-3A yielded two distinct zircon textural and age
The oldest population consists of dark, higher average U cores (>100 ppm) with a weighted average age of 86.6 ± 0.2 Ma (MSWD = 2.3; Figs. 10M and 10N) and average temperatures of ∼753 °C. A younger population of bright, luminescent rims has lower average U values (<75 ppm, n = 19) and yielded a weighted average age of 74.1 ± 0.3 Ma (MSWD = 2.2). Sample 060722-3B yielded a single population with a weighted average of 75.4 ± 0.3 Ma (MSWD = 2.8; Figs. 10O and 10P) and average temperatures of ∼700 °C. These results suggest that the tonalite crystallized at ca. 87 Ma and experienced reheating and amphibolite-facies metamorphism during emplacement of granodioritic dikes at 74-75 Ma. The undeformed character of the dikes indicates that D 3 in the Black Belt shear zone at this locality had ended by ca. 74 Ma.
Site 9 in Deer Canyon (Fig. 3A) exposes an intermediate-to high-strain zone inside the shear zone where granodioritic dikes are interfolded with a tonalite host (Fig. 5G). In intermediatestrain zones, a linear fabric (L > S) defined by coarse, aligned hornblende and plagioclase crystals parallels the axial planes of tight folds. In high-strain zones, the granodiorite dikes were deformed into isoclinal folds and transposed parallel to a mylonitic foliation (S 3 ; Fig. 5H). Sample 060822-9A from the tonalitic host yielded a single population of zircons (n = 50) with a weighted average age of 83.0 ± 0.1 Ma (MSWD = 2.2; Figs. 10Q and 10R). Sample 060822-9B from a tightly folded and sheared granodioritic dike yielded zircon cores (n = 37) with a weighted average age of 83.9 ± 0.2 Ma (MSWD = 2.4) and rims (n = 11) with a weighted average of 77.4 ± 0.4 Ma (MSWD = 1.9; Figs. 10S and 10T). Both samples yielded average temperatures of ∼735 °C. These ages indicate that D 3 deformation at the site occurred from ca. 84 Ma through ca. 77 Ma.
Sample 060822-13B is from an undeformed, biotite-bearing granodiorite dike that cuts both S 2 and S 3 in a metasedimentary host at site 13 (Fig. 3A). This undeformed dike gave a weighted average age of 75.5 ± 0.3 Ma (n = 16; MSWD = 2.2; Figs. 10U and 10V), which we interpret to represent the crystallization age of the dike. Average temperatures were ∼700 °C. The results suggest that D 3 in the Cucamonga terrane at the southern end of the Black Belt shear zone had ended by ca. 75 Ma.
Previous work by Marshak (2016) suggests that the SE San Gabriel Mountains have rotated 20° counterclockwise since the early Miocene. A restoration of all D 2 and D 3 structures by this amount (Fig. 11) brings these features into parallelism with Late Cretaceous faults and shear zones mapped elsewhere in the SCB. Schwartz et al. (2025b) provided a palinspastic reconstruction that shows all these structures form a nearly continuous NW-trending belt once they have been restored to their Cretaceous configurations. From SE to NW, this belt includes the Eastern Peninsular Ranges mylonite zone (Simpson, 1984;Morton et al., 2014), the Black Belt and Cucamonga shear zones (this study), the Alamo Mountain-Piru shear zone (Bixler et al., 2025), the Tumamait shear zone and Sawmill fold-thrust belt (Schwartz et al., 2025a), and the Nacimiento fault (Singleton and Cloos, 2013). The parallelism among them across blocks that record different amounts and senses of late Cenozoic rotation provides an important check on the viability of our restoration.
Another means of evaluating the 20° of counterclockwise rotation is provided by the presence of Miocene grabens (Fig. 9). These faults form part of a suite of transtensional structures within and outside the San Gabriel Mountains (Nourse, 2002). The similar orientations of these structures north and south of the Cucamonga thrust fault suggests that rotations significantly greater than 20° can be ruled out. We therefore conclude A B C Figure 11. Structural measurements restored to their Late Cretaceous orientations by removing 20° of counterclockwise rotation indicated by paleomagnetic data (see text for discussion). (A) L 2 and S 2 structures from D 2 reverse shear zones in the Cucamonga terrane. (B) S 2 folded by F 4 in the Cucamonga terrane (CT). (C) S 3 foliations and L 3 mineral lineations that form part of the Black Belt shear zone in the San Antonio terrane (SAT). Figure 6 shows an explanation of symbols.
that 20° of counterclockwise rotation appears reasonable for the SE San Gabriel Mountains. Despite a lack of published evidence for tilting around a horizontal or inclined axis since the Miocene, we evaluated this possibility using several different means. First, the regional dips of structures preserved in the Cucamonga terrane in Cucamonga Canyon generally are steeper than those in Deer Canyon (Fig. 4). This difference is mainly due to the effects of Cenozoic faulting. For example, the transtensional Demens Canyon fault deforms the contact between the San Antonio and Cucamonga terranes in Cucamonga Canyon (Fig. 9) and locally steepens it by up to 40° on its downthrown (southern) side (Fig. 4B). Away from this fault to the north, the regional dip of Cretaceous foliations matches those in Deer Canyon where faulting is minor. This relationship indicates that the effects of Miocene normal faulting are recognizable in the canyons and suggests that structural measurements in Deer Canyon more closely match pre-Miocene orientations.
The possibility of tilting by slip on minor reverse faults also can be evaluated at the southern ends of Cucamonga and Deer canyons. A comparison of structures inside and outside the areas affected by these faults suggests minor tilting (locally a maximum of 20° from horizontal). The southern end of Cucamonga Canyon, which displays more reverse faults than in Deer Canyon (Fig. 9), shows slightly more tilting (Fig. 4). Since the reverse faults are small, this increase in tilting may result mainly from motion on the larger strands in the Cucamonga thrust fault zone. Previous work suggests that this fault zone dips 25° to 43° to the N near the surface (Morton and Matti, 1987;McPhillips and Scharer, 2018) and steepens to 50° at depth (Cramer and Harrington, 1987). Because of this steepening with depth, restoring up to 1 km of reverse slip (a conservative estimate, since much of the motion appears to be oblique) results in minor (≤20°) tilting of its hanging wall. Consequently, on the basis of our observations of faulting in the study area, we conclude that the effects of tilting are recognizable and can be quantified, and that Deer Canyon preserves dip orientations that are closest to their pre-Miocene configurations. These relationships support the interpretation that the canyons expose a tilted crustal section with deeper levels exposed to the south next to the Cucamonga thrust.
A final approach to evaluating the effects of tilting is to assess its potential effects on our kinematic analyses of Cretaceous structures. One advantage of the approach we used is that the assessment of the regional kinematics of D 2 and D 3 deformation is insensitive to reorientation by younger faulting. For example, the interpretation that the F 4 domes and basins are compatible with horizontal shortening and vertical thickening is independent of the regional dip of foliations and does not rely on the precise knowledge of pre-faulting orientations. Even if the Black Belt and Cucamonga shear zones originally were horizontal (e.g., detachments) and later were steepened by slip on Cenozoic faults, the geometry of the domes and basins still allows us to determine that they formed by crustal shortening and thickening. This pattern suggests that even if the orientations of structures in Deer Canyon were tilted significantly, we can still distinguish Late Cretaceous crustal shortening and thickening from younger extension and crustal thinning.
The structural data reported here define a major lower-crustal shortening and thickening event that affected the SCB prior to the Late Cretaceous emplacement of the Tonalites of San Sevaine Lookout. This event (D 2 ) is represented by a 2.5-km-thick (minimum) package of S 2 foliations that record a top-to-the-SSE (restored) reverse sense of shear parallel to a down-dip hornblende-plagioclase mineral lineation (L 2 ; Fig. 11A). The intensity of subsequent phases of deformation (D 3 , D 4 ) and metamorphism (M 2 ) make it difficult to establish the absolute timing of D 2 , and its precursor (D 1 ), with certainty. Nevertheless, three observations provide new limits on its age and duration.
First, because D 2 only occurs in the Cucamonga terrane and not in the San Antonio terrane, the age of the oldest tonalite in the latter provides a limit on how young D 2 can be. The oldest tonalite we found is in Deer Canyon (sample 060722-3A) and crystallized at ca. 87 Ma. Schwartz et al. (2023Schwartz et al. ( , 2025a) ) reported similar ages from tonalites in the same region beginning at ca. 93 Ma, indicating that D 2 occurred prior to that time. Second, we report textural data showing that D 2 occurred simultaneously with granulite-facies metamorphism and migmatization in the Cucamonga terrane (e.g., Fig. 5B). Zircon data published by Schwartz et al. (2025a) showed that high-temperature metamorphism and partial melting of these rocks (collectively referred to as M 2 in this study) accompanied arc magmatism at ca. 124 Ma (their sample 28b) and ca. 94 Ma (their sample 25D). Third, the zircon ages and geochemistry of samples 21CM51b-H and -V1 (this study) provide a limit of 91-86 Ma for how young D 2 can be, which is similar to the limit provided by the age of the oldest tonalites. We therefore conclude that D 2 most likely occurred during the 124-93 Ma interval, placing it within the time frame of the Sevier fold and thrust belt (Fig. 12A). The timing and significance of the D 1 event remain unclear.
During the Early to mid-Cretaceous, as the rate of magmatism in the arc was increasing (Cecil et al., 2012;Paterson and Ducea, 2015;Schwartz et al., 2023), several other regional tectonic events occurred that help us interpret the origin and significance of the D 2 event. Throughout the 125-90 Ma interval, convergence rates between the Farallon and North American plates increased and the orogenic front propagated eastward into the western Sevier fold-thrust belt (Yonkee et al., 2019;Yonkee and Weil, 2015). This latter belt is a thin-skinned, foreland-propagating (west to east) wedge that accumulated up to 220 km of shortening from Cretaceous to Paleogene time (DeCelles and Coogan, 2006;Yonkee and Weil, 2015). The shallow (10-12 km depth), thin-skinned shortening associated with this belt is thought to have been balanced by lower crustal thickening and the uplift of an orogenic plateau in the hinterland to the west (DeCelles and Coogan, 2006;Yonkee et al., 2019;Yonkee and Weil, 2015). The thick-skinned style of shortening we describe as D 2 appears to form part of this postulated system of lower crustal thickening (Fig. 12A). These relationships suggest that D 2 formed part of an orogenic system where plate convergence rates, arc magmatism, lower-crustal shortening in the hinterland, and thin-skinned shortening in the foreland all were dynamically linked.
New U-Pb zircon ages show that the Black Belt and Cucamonga shear zones formed during a surge in arc magmatism in the SCB at 84-74 Ma. The two shear zones record identical ages, similar monoclinic strain symmetries (Figs. 7A and 7B), and matching top-to-the-SW (restored) sinistral-reverse (arc oblique) displacements. This deformation (D 3 ) also overlapped in time with an episode of folding (F 4 ) that created a regional dome-and-basin interference pattern (Fig. 8). Together, these structures created a thick-skinned, partitioned style of intra-arc transpression where arc-parallel displacements were accommodated mainly by simple-shear dominated D 3 shear zones deforming by monoclinic strain symmetry, and a large component of arc-normal shortening was accommodated by the F 4 domes and basins. Although the total amount of shortening appears small (less than a few kilometers), the onset of transpression by ca. 84 Ma (Fig. 12B) marks an important shift away from the distributed style and purely reverse displacements that characterized the older D 2 event (Fig. 12A).
Other structures in the SCB record similar styles of transpression that overlapped in time with the structures we report from Cucamonga and Deer canyons. In the Mt. Pinos region of the Transverse Ranges, the Tumamait shear zone accommodated sinistral-reverse (top-up-to-the-SW) displacements from at least 77 Ma through 70 Ma (Fig. 12B; Schwartz et al., 2025a). Like the Black Belt and Cucamonga shear zones, this structure is thin (0.5 km thick), mylonitic, and formed at high temperatures (699-718 °C) during the emplacement of sheeted plutons. It also is deformed by late folds that were kinematically linked to NE-directed motion on the Sawmill thrust fault (ST, Fig. 12C; Schwartz et al., 2025a). This latter thrust forms the boundary between the SCB and underthrust trench sediments of the Pelona schist and is interpreted to have been active until at least ca. 66 Ma (Schwartz et al., 2025a). Other mid-crustal shear zones that record sinistral or sinistralreverse transpression include the Nacimiento fault at ca. 72 Ma (Singleton and Cloos, 2013), the Alamo Mountain-Piru Creek shear zone at 76-72 Ma (Zucker, 1990;Bixler et al., 2025), and the La Paz fault in southern Baja California at 84-56 Ma (Mattern et al., 2010). These structures define a transpressional fold-thrust belt that shortened and imbricated the middle and lower crust of the Late Cretaceous Southern California arc from ca. 84 Ma until at least ca. 72 Ma (Fig. 12B).
East of the SCB, in what is now west-central Arizona and southernmost Nevada, episodes of mid-to Late Cretaceous thrusting formed the Maria fold-thrust belt (Reynolds et al., 1986;Laubach et al., 1989;Boettcher et al., 2002;Strickland et al., 2018;Wong et al., 2023). This belt records shortening at high angles to the Cretaceous arc without significant lateral displacements. Boettcher et al. (2002) reported that, in at least some parts of the belt, the shortening occurred during and after the emplacement of 86-78 Ma granites. Cawood et al. (2022) interpreted a younger age (68-63 Ma) for thrusting in the Cargo Muchacho Mountains in the southernmost part of the belt. This deformation, thus, overlapped in time with intra-arc transpression and helped create a thick lower-crustal welt east of and below the arc that may have reached thickness up to 57 ± 12 km (Fig. 12B; Chap- , 2020). Since no significant arc-parallel displacements have been reported east of the SCB, it appears that lateral (arc-parallel) displacements preferentially were partitioned into the magmatic arc at this time. Similar patterns of deformation partitioning occur in other obliquely convergent margins where arc-parallel motion tends to focus into the magmatic arc where the crust has been weakened by magma and heat (Tikoff and de Saint Blanquat, 1997;McCaffrey et al., 2000;Corti et al., 2005;Klepeis et al., 2022).
By ca. 70 Ma, sinistral transpression within the arc had transitioned to folding and thrusting as the angle of subduction between the Farallon and North American plates shallowed (Fig. 12C). This shallowing resulted in the underplating of trench and accretionary wedge sediments beneath both the SCB (Schwartz et al., 2025a) and the Maria fold-and-thrust belt after 75-70 Ma (Jacobson et al., 2011;Haxel et al., 2014;Strickland et al., 2018). The sediments now appear as the Pelona and Orocopia schists (Fig. 1; Jacobson et al., 2017;Seymour et al., 2018). Cawood et al. (2022) and Schwartz et al. (2023Schwartz et al. ( , 2025b) attributed these events to flat-slab subduction of the oceanic Hess Plateau beneath North America at 70-66 Ma. Flat-subduction also has been linked to the end of arc magmatism and a regional cooling of the arc to below 350 °C (Schwartz et al., 2023). Schwartz et al. (2025b) review the timing and locations of some of the other ductile thrust systems that accommodated subduction-related convergence during this period.
The predominance of sinistral and sinistralreverse displacements within the SCB during the Late Cretaceous contrasts with the mostly dextral sense of shear recorded in other intra-arc transpressional shear zones located north of the present-day Garlock fault (Fig. 1; Tikoff et al., 2023). To explain the opposing shear senses, as well as differences in the timing and style of arc magmatism, Schwartz et al. (2025b) postulated a tectonic boundary between the two arc segments (see also Umhoefer and Blakey, 2006). These authors suggested that the Kula-Farallon-North America plate triple junction was located at the northern end of the SCB at ca. 85 Ma thereby segmenting the California arc at the presentday location of Garlock fault. This hypothesis is based on work by Umhoefer (1987) and Umhoefer et al. (1989) who proposed that the triple junction formed at the southern edge of the Insular superterrane between the Peninsular Ranges batholith and the SCB during a major plate reorganization at 90-85 Ma (see also Thorkelson and Taylor, 1989;Umhoefer, 2019). Its location near the present-day Garlock fault explains the regional continuity of the sinistral-reverse-style of intra-arc transpression within the SCB and southward from 85 Ma to 70 Ma.
The development of sinistral transpressional shear zones (D 3 ) at ca. 84 Ma immediately followed the start of a Late Cretaceous magmatic surge at 94-86 Ma and lasted until at least ca. 74 Ma (Schwartz et al., 2023(Schwartz et al., , 2025a(Schwartz et al., , 2025b, this study), this study). We found that granulite-facies metamorphism (M 2 ) accompanied this magmatism and is recorded in lower crustal gneisses of the Cucamonga terrane at Cucamonga Canyon (site 51b, Fig. 3B). This event coincided with the initial emplacement of a large mass of tonalite into the middle crust at ca. 92-86 Ma. Schwartz et al. (2025b) distinguished this same event at 94-87 Ma in Deer Canyon (their samples 25A-25D). Cooling and a transition to lower temperature, upper amphibolite-facies metamorphism occurred from ca. 77 Ma to at least ca. 74 Ma and coincided with both D 3 deformation and the emplacement of granodiorite dikes throughout the middle and lower crust (see also Schwartz et al., 2025b). These results show that while the Black Belt and Cucamonga shear zones were forming, the middle and lower crust remained hot, weak, and partially molten, with temperatures reaching 749-780 °C during the 84-74 Ma interval and probably until ca. 72-70 Ma.
Structural relationships between D 3 fabrics and sheeted dikes in the San Antonio terrane illustrate the influence of magmatism on the architecture of the Black Belt shear zone. Variations in grain size, composition, temperature, and melt concentration created rheological contrasts that localized D 3 deformation at dike margins, forming widely spaced high-strain zones within a 1-km-thick zone of deformation (Fig. 3). The effects of rheological contrasts created by intruding magma also is evident in the asymmetry of the shear zone. Most of the deformation localized within the San Antonio terrane rather than in the Cucamonga terrane, showing that, at the time the shear zone formed, the mid-crustal San Antonio terrane was weaker than the lowercrustal gneisses of the Cucamonga terrane.
Differences in the internal structures of the Black Belt and Cucamonga shear zones also highlight the influence of syn-tectonic magma emplacement in the former example. In general, the latter shear zone exhibits lower volumes of syn-tectonic and post-tectonic dikes, less mylonite, more penetrative gneissic and schistose fabrics, and a greater abundance of disharmonic F 4 folds than the former. An analysis of the L 3 /S 3 fabric in areas of low-, intermediate-, and highstrain zones in the Cucamonga shear zone shows that its evolution was controlled mainly by the reactivation of preexisting S 2 foliations (e.g., Fig. 7A) rather than magma emplacement. Similar influences of preexisting structures on shear zone initiation and growth are well-documented in other transpressional systems worldwide, including in Fiordland, New Zealand (Buriticá et al., 2019;Klepeis et al., 2022;Miranda et al., 2023b); SE Asia (Li et al., 2022); the western United States (Lee et al., 2012); Scotland (Armitage et al., 2021); and elsewhere.
Relationships between magmatism and deformation similar to the ones we report also have been described for other regions. In southwest New Zealand, for example, Klepeis et al. (2022) showed that episodic magma emplacement influenced shear zone architecture and drove its growth. As magma entered the lower crustal roots of a steep transpressional shear zone and began to solidify, strain hardening helped drive deformation out of the strengthening lower crust and into a weaker middle crust where the system widened by a factor of three and connected to fold-thrust belts. The Black Belt shear zone also shows that the episodic intrusion of dikes whose physical properties change over time was a primary driver of geometric complexity, including the formation of widely spaced high-strain zones. Strain hardening as a result of magma solidification also appears to have led to a widening of the zone of deformation, resulting in the formation of late F 4 folds as shortening continued and localized shearing in highstrain zones ceased. Comparable relationships have been observed in the western Idaho shear zone (Giorgis et al., 2005), the Cascades (Miller et al., 2009), and the southern Sierra Nevada batholith (Saleeby et al., 2008). These studies suggest that magma emplacement and solidification influences shear zone width, increases the degree of coupling between intruding magmas and their host rock, and influences displacement partitioning.
Interlayered pseudotachylyte and mylonite similar to the example we describe (Figs. 5J and 5K) is thought to represent a cyclic interplay between seismic events and ductile viscous creep (mylonitization) caused by ductile instabilities (White, 2012;Stewart and Miranda, 2017) and/or fluid infiltration (Menegon et al., 2017;Jamtveit et al., 2019;Hawemann et al., 2019;Wex et al., 2019). Michalchuk et al. (2023), for example, showed that rheological weakening of the lower crust leading to viscous creep can occur when a fluid infiltrates a transiently permeable (fractured) shear zone, thereby facilitating diffusive mass transfer and creep. However, the formation of coeval mylonite and pseudotachylyte is not well understood and remains intensely debated (see review by Menegon et al., 2021;White, 2012).
In the Black Belt shear zone, alternating layers of brittle and ductile deformation form an integral part of high-strain zones whose spacing and evolution were controlled by magma emplacement and solidification. The textures we describe show that strain localization began during a rheological transition when viscous flow in crystal-rich magmas gave way to both faulting and crystal-plastic deformation as individual dikes solidified. Pseudotachylyte generation during this transition was followed by renewed viscous creep that formed mylonite, marking a change in both strain rates and crustal strength. We note that in the Black Belt shear zone, pseudotachylyte formed in rocks rich in hydrous phases (biotite, hornblende) and are conspicuously absent in lower-crustal granulites of Cucamonga terrane. This latter observation contrasts with observations in other belts where pseudotachylyte is thought to reflect deformation of dry strong, anhydrous rocks typical of lower-crustal granulites (Menegon et al., 2021).
These observations tentatively support a view where strain localization and high strain rates occurred in dikes when enough solid material had crystallized to allow ductile rupture. We speculate that the high ambient temperatures indicated by episodes of amphibolite-and granulite-facies metamorphism, and residual fluids in the crystallizing magmas, may have helped promote high strain rates. Despite these intriguing possibilities, more work needs to be done to determine exactly how these fabrics formed, including which mineral phases contributed to melting and whether hydrous or anhydrous conditions prevailed (Miranda et al., 2023a).
The southeastern San Gabriel Mountains record a polyphase history of magmatism, hightemperature metamorphism, and deformation that illustrates how strain localized deep within the mid-lower crustal root of the SCB during Late Cretaceous orogenesis. South-directed (restored orientation) ductile thrusting and lower-crustal thickening at 124-93 Ma accompanied uplift of an orogenic plateau in the hinterland and thin-skinned crustal shortening in the Sevier fold-thrust belt as an orogenic front propagated eastward. The mid-crustal Black Belt shear zone and lower-crustal Cucamonga shear zone formed together at ca. 84 Ma shortly after a flare-up in arc magmatism began. Both shear zones record sinistral-reverse displacements from ca. 84 Ma until at least ca. 74 Ma and probably until ca. 72-70 Ma. Each forms part of a thick-skinned, kinematically partitioned style of intra-arc transpression where arc-parallel and arc-oblique displacements were accommodated on dipping surfaces within the two shear zones and a component of arc-normal shortening was accommodated by coeval domes and basins. This deep-crustal transpressional fold-thrust belt is correlative with other Late Cretaceous transpressional faults and shear zones mapped throughout the SCB.
The internal architecture and evolution of the Black Belt shear zone was strongly influenced by magmatism. Rheological contrasts between comagmatic intrusions localized deformation into a series of thin (10-30 m), widely spaced highstrain zones whose geometry reflects an igneous architecture created by sheeted, episodically emplaced dikes. Shear zone fabrics record deformation that initiated during hypersolidus flow in mingling tonalite, diorite, and granodiorite bodies and continued to develop as the rocks transitioned below the solidus where crystal-plastic processes produced mylonitic foliations. As the intrusions solidified, strain hardening increased shear zone width, increased the degree of coupling between dikes and their host rock, and increased displacement partitioning. Syn-magmatic high-strain zones record alternations between ductile rupture and brittle fracturing that produced pseudotachylyte and viscous creep that formed mylonite. This interlayering appears to reflect transient seismic events that occurred once the crystallizing magmas had solidified and strengthened.
A structural analysis suggests that 20° of Cenozoic counterclockwise rotation is reasonable for the SE San Gabriel Mountains. The effects of horizontal tilting is spatially variable and linked to motion on early Miocene normal faults and Pliocene-Quaternary reverse faults. Deer Canyon preserves structural orientations that are closest to their pre-Miocene configuration.
Geological Society of America Bulletin, v. 136, no. XX/XX
Geological Society of America Bulletin, v. 136, no. XX/XX