Structure, metamorphism, and geodynamic significance of the Catalina Schist terrane
Published:May 18, 2020
- PDF LinkChapter PDF
John P. Platt*, Marty Grove*, David L. Kimbrough*, Carl E. Jacobson*, 2020. "Structure, metamorphism, and geodynamic significance of the Catalina Schist terrane", From the Islands to the Mountains: A 2020 View of Geologic Excursions in Southern California, Richard V. Heermance, Joshua J. Schwartz
Download citation file:
This guide begins with an overview of the internal structure and petrology of the Catalina Schist terrane as exposed on Santa Catalina Island, California, followed by a discussion of the tectonic setting and exhumational history of the terrane, and the Cenozoic tectonic and geological evolution of the Inner Borderland, within which it lies. The guide then presents an itinerary for a three-day field trip from 9–11 May 2020. Next, we present a tectonic model for the formation of the Catalina Schist, followed by a discussion of its relationship to the Pelona, Orocopia, Rand, and related schists in southern California.
THE CATALINA SCHIST TERRANE
Santa Catalina Island is one of several islands offshore southern California, and is unique in that it exposes a metamorphic complex including blueschist to amphibolite facies rocks, generally known as the Catalina Schist. The only other on-land exposure of these rocks is on the Palos Verdes peninsula south of Los Angeles, but drill-hole data, offshore dredging, and the distribution of Catalina Schist clasts in the Miocene San Onofre Breccia suggest that the terrane underlies much of what is known as the Inner Borderland, extending from Santa Cruz Island in the north at least as far south as San Diego (Howell and Vedder, 1981; Legg, 1991).
The metamorphic rocks on Santa Catalina Island were first described as the Catalina Schist by Woodford (1925), who suggested that they were related to the Franciscan Complex in the Coast Ranges of California. Since then, it has been recognized that there are several distinct elements within the Catalina Schist accretionary complex that differ in age, metamorphic evolution, and structural history (Bailey, 1941; Platt, 1975; Grove and Bebout, 1995). These rock packages formed at different locations within the accretionary complex and were ultimately juxtaposed during the mid-Cretaceous (Grove et al., 2008). The accreted rocks now form a series of regionally gently dipping thrust sheets. They are described here for convenience from the bottom up.
Catalina Blueschist Unit
The Catalina Blueschist Unit crops out at low structural levels in the central part of the island (Figs. 1.1–1.3), and in this account includes the entire west end of the island beyond the isthmus. In many respects, this body resembles Franciscan rocks of Cretaceous age in the Diablo Range and the Eastern belt of the Northern Coast Ranges. It comprises metamorphosed graywacke sandstones and shales, which make up ~80% of the outcrop area, together with metamorphosed conglomerates, chert, volcanic rocks mainly of basaltic composition, and irregular bodies of serpentinite and related rocks. The geometric occurrence of these rocks ranges from relatively coherent, meter to kilometer scale, mono-lithologic domains that are disrupted by highly sheared, compositionally heterogeneous, and variably metasomatized domains that are crudely isofacial in metamorphic grade to the rocks that they separate.
Metamorphic grade in the central part of the island is predominantly in the lawsonite-blueschist facies, with widespread sodic amphibole and lawsonite, sporadic jadeitic pyroxene, some omphacite, and rare garnet. Estimated metamorphic conditions (Fig. 1.4) lie the range 300–400 °C and 8–11 kbar (Sorensen, 1987). Lower temperature and less devolatilized rocks on the NW tip of the island largely lack sodic amphibole, however, and range from lawsonite-albite facies to low greenschist facies with actinolite + epidote (Bebout and Fogel, 1992; Bebout et al., 1993; Grove and Bebout, 1995). It is uncertain whether these differences reflect a transitional decrease of metamorphic grade, or presently unrecognized tectonic boundaries.
Detrital zircon ages from lawsonite-albite through lawsonite blueschist facies metagraywackes of the Catalina Blueschist Unit indicate a maximum depositional age of 97 ± 3 Ma (Grove et al., 2008). The measured U-Pb age distribution strongly resembles both the emplacement history of the adjacent 130–90 Ma Peninsular Ranges batholith and derivative Cenomanian forearc sediments (Grove et al., 2008). The protolith ages of the mafic rocks and chert are unconstrained, however, and could be significantly older. 40Ar/39Ar incremental heating performed with phengite yields strong age gradients down to 20 Ma and terminal ages between 100 and 90 Ma (Grove and Bebout, 1995). Taking both the detrital zircon U-Pb and 40Ar/39Ar cooling ages into consideration, it seems most plausible that the Catalina Blueschist Unit formed in mid-Cretaceous time (100–90 Ma) and cooled slowly at depth throughout the Late Cretaceous–Early Cenozoic. Finer-grained Cr-rich fuchsites from highly sheared mixtures of sedimentary and ultramafic rocks associated with the lawsonite-albite facies metasediments yield similar age gradients, with terminal ages between 85 and 60 Ma that further support slow cooling throughout the Late Cretaceous and Early Cenozoic.
The structural style in the Catalina Blueschist Unit is complicated and difficult to analyze. On the north side of the island, and on the NW end, there appear to be regular intercalations of metabasalt-metachert-metagraywacke sequences that are suggestive of thrust slices formed during subduction and accretion. Metagraywackes variably preserve bedding and detrital textures (Fig 1.5A); detrital feldspar is commonly replaced by jadeitic pyroxene or lawsonite (Fig. 1.6A). Metaconglomerate (Fig. 1.5B) preserves volcanic, dioritic, and gabbroic clast assemblages that are highly metasomatized (Sievers et al., 2017). Features present within coherent metasedimentary rocks range from massive beds to tightly folded layered domains that exhibit axial planar pressure-solution cleavage and/or more brittle crack-seal relationships involving abundant quartz veins. Metavolcanic rocks include undeformed pillow basalt, pillow breccia, and diabase, some of which preserve relict igneous textures and minerals. These less deformed mafic rocks are characteristically enveloped by highly deformed and metasomatized glaucophane-lawsonite schist (Figs. 1.6B, 1.6C). There is evidence for polyphase deformation in schistose metavolcanic and metasedimentary rocks in the form of refolded folds, crenulation cleavage, and flattened and folded clasts in fragmental rocks.
Elsewhere, the Blueschist Unit shows a much more disruptive style of deformation, involving extensive boudinage, accompanied by abundant quartz veining. In places, the rocks are compositionally very heterogeneous on centimeter to meter scales, and consist of mixtures of strongly deformed sandstone, shale, and mafic and ultramafic material that are crudely isofacial with adjacent coherent lithologies (Fig. 1.5C). It is uncertain whether these sheared domains represent tectonic mélange and/or shearing of conglomerates formed from sedimentary, volcanic, and ultramafic clasts (e.g., Wakabayashi and Dilek, 2011).
Sorensen (1987) suggested that the mélanges of the Blueschist Unit contain tectonic blocks of higher grade than their surroundings, including high-grade garnet blueschist and possible eclogite. An ~150 m lens of relatively coarse-grained, garnet-bearing lawsonite blueschist surrounded by metaconglomerate and mélange-type rocks crops out south of Little Harbor. Phengite from this block yields a total gas 40Ar/39Ar and Rb/Sr age of ca. 135 Ma (Grove et al., 2008). However, the history experienced by the block is far more complex than this concordance implies. Age gradients yielded by phengite from the sample and another like it rise up to 160 to 150 Ma (Grove and Bebout, 1995). The presence of garnet and rutile suggests that these rocks reached a higher metamorphic temperature than the surrounding blueschists. These observations suggest that the garnet-bearing blueschists have an older metamorphic history than the bulk of the mid-Cretaceous Blueschist Unit.
A distinctive feature of the Blueschist Unit is the abundance of veins, which include syn-metamorphic veins with quartz, sodic amphibole, lawsonite, and rare aragonite, and post-metamorphic veins with albite, quartz, calcite, and pumpellyite (Bebout and Barton, 1989). These veins commonly show crack-seal textures, quartz fibers, and evidence for fluid-filled cavities. In places there are closely spaced sets of sub-parallel quartz veins, which may be tightly folded (Fig. 1.5D). These “sheeted vein complexes” may mark zones of hydraulic fracturing accompanying thrusting within the unit, and form part of a suite of structural features that may have been a source of the phenomena known as seismic tremor and slow slip, which are observed on present-day active margins such as the Cascadia subduction zone (e.g., Peng and Gomberg, 2010; Schwartz and Rokosky, 2007; Burgmann, 2018).
Catalina Greenschist Unit
The Blueschist Unit is overlain structurally by thrust slices of rock a few hundred meters thick, with a similar protolith assemblage to the Blueschist Unit, but distinctly higher metamorphic temperatures. In this guide, we use the term “Catalina Greenschist Unit” after Platt (1976) as a single map unit for these slices, while accepting that they vary in metamorphic grade from epidote blueschist facies to albite epidote amphibolite facies, and exhibit other significant differences in age and tectonic history (Jacobson and Sorenson, 1986; Grove and Bebout, 1995; Grove et al., 2008). All rocks included in the Greenschist Unit have epidote/clinozoisite as the Ca-Al-silicate phase, in place of lawsonite in the Blueschist Unit. Another distinctive feature of all Greenschist Unit metasedimentary rocks is the abundance of porphyroblasts of albite (e.g., black-albite schist, Fig. 1.6E). The albite porphyroblasts characteristically appear dark due to a high concentration of included graphite.
Biotite and garnet are common in comparatively coarse-grained epidote amphibolite facies mica schists derived from graywacke sandstones and shales. Quartz schists with piemontite (Mn-epidote), riebeckitic amphibole, garnet, and stilpnomelane were formed from original chert layers (Figs. 1.5F, 1.6F, 1.6G). Metamafic rocks with either sodic or calcic amphibole are derived from metabasalts that locally preserve pillow textures (Fig. 1.5E). Fine-grained mafic schists present within epidote blueschist facies rocks characteristically contain 5–10-mm-long epidote porphyroblasts and alternating layers enriched in either sodic or calcic amphibole (Fig. 1.6D). Metamorphic conditions in the various component slices of the Greenschist Unit are poorly constrained (Fig. 1.4), but likely lay in the range 450–580 °C, with pressures in the range 7–12 kbar (Sorensen, 1987). Some of the slices appear to show some internal variation in metamorphic grade, but none of the mapped exposures contain the full range of grade. The wide variation in mineral assemblages may partly reflect changing conditions during the metamorphic history.
Detrital zircon ages from an epidote blueschist facies slice in Cottonwood Canyon indicate a maximum depositional age of 100 ± 3 Ma, and an overall age distribution that strongly resembles the lawsonite-albite and lower-grade rocks (e.g., Peninsular Ranges batholith provenance signature; Grove et al., 2008). Phengite 40Ar/39Ar ages also overlap with those from the lawsonite blueschist facies rocks (99–95 Ma). In contrast, an albite-epidote amphibolite slice in Little Springs Canyon/Little Harbor) indicates a considerably older maximum depositional age of 113 ± 3 Ma. The provenance signature of the higher-grade rocks is also distinct in that it contains appreciable Jurassic and older Proterozoic zircon characteristic of a more cratonal provenance (Grove et al., 2008). Coarse phengite yields relatively flat 40Ar/39Ar age spectra and total gas ages between 102 and 97 Ma.
The internal structure of the Greenschist Unit is complicated but fairly coherent. Pervasive schistosity reflects ductile deformation to large strains, and this schistosity is commonly microfolded, producing a crenulation cleavage. Polyphase folding has locally produced fold interference structures, but correlation of individual phases of folding is difficult. Metachert layers form good markers that can be traced for distances of up to 3 km, and are folded on various scales (Figs. 1.3, 1.7, 1.8).
Contact relationships with the other major map units in the Catalina Schist generally involve low-angle faults. Metachert near the boundary with the overlying Amphibolite Unit shows mylonitic microstructure with dynamic recrystallization of the quartz. Conversely, a highly distinctive feature of the thrust contact with the underlying Blueschist Unit is an ultramafic mélange zone up to 100 m thick. This strongly sheared and metasomatized zone contains a variety of tectonic blocks that include (1) garnet amphibolite and less common eclogite with clinopyroxene containing 20%–35% jadeite component recrystallized to garnet amphibolite (Fig. 1.5H); (2) predominantly coarse-grained garnet-hornblendites; and (3) lenses of serpentinite (Platt, 1976). Much of this material shows strong affinities to the Amphibolite Unit and the ultramafic mélange that overlies it (see following section).
Catalina Amphibolite Unit
The Amphibolite Unit of the Catalina Schist is a sheet of rock several hundred meters thick and ~6 km in lateral extent that overlies both the Greenschist and Blueschist Units. The bulk of the unit is made up of mafic amphibolite (Figs. 1.5G, 1.6H) with pale-green pargasitic hornblende, clinopyroxene, zoisite/clinozoisite, and intermediate plagioclase. The plagioclase is almost entirely altered to epidote group minerals, white mica, and sodic plagioclase. The amphibolite is weakly layered in terms of mineral proportions, and is remarkably homogeneous. Layering is much steeper than the tectonic boundaries of the unit and accounts for an ~2000 m thickness measured normal to compositional layering. The layering, its overall homogeneity, and its Mg-rich bulk composition (Ernst, 1970) suggest that it may have originally formed part of a layered gabbroic intrusion. No trace of any original igneous texture is preserved, however. The mafic amphibolite is locally cut by pegmatitic dikes composed of varying proportions of hornblende, plagioclase, quartz, and muscovite, which appear to reflect partial melting during metamorphism (Sorensen, 1988b).
The mafic amphibolite is overlain in places by bodies of metasedimentary rock. The most important type volumetrically is a migmatitic paragneiss. The protolith was a siliciclastic sediment, but it is now a coarse-grained metamorphic rock with quartz, plagioclase, muscovite, biotite, garnet, zoisite, and kyanite, together with irregular segregations, layers and dikes of tonalitic material, locally pegmatitic. These segregations appear to have formed by partial melting during high-grade metamorphism. The paragneiss also locally contains interlayers of garnet amphibolite, presumably originally of igneous origin.
Coarse-grained massive quartzite, with irregular layers and patches of fine-grained Mn-rich garnet, likely represents metachert. A thin but continuous layer several hundred meters in lateral extent directly overlies the mafic amphibolite west of Cottonwood Canyon, and it also occurs interlayered with the paragneiss.
Detrital zircon ages from the migmatitic metasedimentary rocks of the Amphibolite Unit indicate a maximum depositional age for the protolith sediment of 122 ± 3 Ma (Grove et al., 2008). Most detrital zircons are overgrown by thick and irregularly formed rims of metamorphic zircon that define a poorly defined maximum at ca. 112 Ma. The overall detrital age distributions include significant Jurassic and Proterozoic zircon that account for between 20% and 50% of the total measurements. This cratonal provenance is significantly different from the Blueschist Unit, which is dominated by zircons of magmatic arc provenance. Protolith ages for other components of the unit are unknown, but the timing of metamorphism and cooling has been constrained by a number of techniques. Lu-Hf dating of garnet from a mafic interlayer in the migmatitic paragneiss indicates peak metamorphism at 114.5 Ma (Anczkiewicz et al., 2004). The K-poor hornblendes from the coherent mafic amphibolite yield poorly defined total gas ages between 110 and 100 Ma (Grove and Bebout, 1995), and muscovites from the overlying metasediments yield total gas ages between 105 and 100 Ma. The 10 m.y. contrast between the Ar ages and the U-Pb and Lu-Hf metamorphic ages indicates slow cooling after the attainment of amphibolite facies conditions.
The coherent amphibolite facies rocks are structurally overlain by a large body of ultramafic-dominated mélange (Figs. 1.3, 1.7). This assemblage poses some of the most difficult questions of interpretation in the entire complex. In terms of whole rock composition, the mélange appears to be a 2:1 mixture of ultramafic to mafic protoliths (Bebout and Barton, 2002). Ultramafic domains include massive serpentinized peridotite, with a relict fabric defined by oriented enstatite. However, much of the original ultramafic rock has been pervasively metasomatized and strongly deformed to form Si-rich schists that contain talc, serpentine, Mg-chlorite, Mg-amphibole (cummingtonite and anthophyllite), and Ca-amphibole (tremolite-actinolite) (Sorensen, 1988b). Radial sprays (e.g., rosettes) up to 30 cm across of pyroxene, amphibole, and talc in metasomatized ultramafic blocks mantled by the schists indicate that high-temperature metamorphic recrystallization preceded deformation.
Mafic-rich domains contain abundant blocks of amphibolite facies rock, most notably coarse-grained rocks rich in dark Fe-Ti–rich hornblende and garnet. Less common blocks consist of plagioclase-bearing garnet amphibolites with migmatitic fabrics and garnet-bearing quartzite that appears to be derived from the underlying coherent Amphibolite Unit. These blocks vary in dimension from several cm to several hundred meters and lie within a coarse-grained matrix of chlorite-amphibole schist (Fig. 1.9). Pegmatitic quartz + plagioclase lenses occur sporadically throughout the chlorite-rich mélange. Sorensen and Barton (1987) estimated conditions of 8–11 kbar and 640–750 °C for metamorphism and partial melting in the Amphibolite Unit (Fig. 1.4), based on mineral assemblages (in particular the assemblage anthophyllite + talc), garnet-clinopyroxene thermometry, metamorphic reactions calculated in the system CaO-Al2O3-TiO2-SiO2-H2O, and melting relations in the granite system. More recent estimates from Penniston-Dorland et al. (2018) that are based upon Zr-in-rutile thermometry from the garnet-hornblendite blocks range from 650 to 730 °C. The presence of eclogitic patches and layers (Fig. 1.5F) in some of the tectonic blocks, containing clinopyroxene with 12%–22% jadeite (Sorensen and Barton, 1987), may suggest an earlier somewhat higher-pressure or lower-temperature history.
Mattinson (1986) obtained U-Pb ages of 112 and 114 Ma on titanite from garnet hornblendite blocks in the ultramafic mélange, and Page et al. (2019) report a U-Pb age on metamorphic zircon rims from a garnet quartzite block of 115.1 ± 2.5 Ma. 40Ar/39Ar ages on hornblende from garnet amphibolite blocks lie in the range 116–108 Ma (Grove and Bebout, 1995). These results compare well with those previously mentioned from the coherent portion of the Amphibolite Unit, and provide further evidence for slow cooling of all amphibolite facies rocks between 115 and 100 Ma.
Metasomatism in the Catalina Schist
All the tectonic units that make up the Catalina Schist show evidence for extensive metasomatism, in the form of veins, reaction rinds on blocks in mélange, and changes in bulk chemistry of many rock types, most notably the ultramafic rocks (Bebout and Barton, 1989, 1993, 2002; Bebout et al., 1999). Stable isotope and whole-rock geochemical evidence indicates infiltration of the higher-grade units with a relatively homogeneous fluid (Bebout, 1991). Systematic measurements of δ11B, δ13C, δ15N, and δ18O throughout the Catalina Schist suggest that the most likely fluid source originated from devolatilization of low-grade sedimentary rocks, equivalent to those that make up the bulk of the Blueschist Unit (Bebout, 1991; Bebout and Fogel, 1992; Bebout et al., 1993; Bebout and Barton, 2002).
Metasomatic assemblages in each of the units of the Catalina Schist tend to be crudely isofacial with mineral assemblages in less altered lithologies. For example, blocks of lawsonite blueschist incorporated into sheared domains within the Blueschist Unit show reaction rinds of actinolite, chlorite, phengite, and albite. These rinds are enriched in highly siderophile elements (HSE, including Os, Ir, Ru, Pt, Pd, and Re) and MgO, Cr, and Ni, suggesting interaction with an ultramafic source (Penniston-Dorland et al., 2010). Observations on veins, mélange matrix, and cobbles in these rocks suggest mobilization of Li during high pressure-temperature (P-T) fluid–rock interactions, which also resulted in enrichments in B, N, and some large ion lithophile elements (K, Rb, Cs, Ba) (Penniston-Dorland et al., 2010; Sievers et al., 2017). Ultramafic lithologies present within the Greenschist Unit are pervasively metasomatized to talc-actinolite schists. Within the ultramafic mélange of the Amphibolite Unit, serpentinite blocks have rinds composed of magnesiohornblende + anthophyllite + talc + chlorite + quartz + rutile (Sorensen and Barton, 1987). Alternatively, garnet amphibolite blocks are mantled by a more aluminous metasomatic assemblage consisting primarily of chlorite and actinolitic amphibole. In both cases, the mélange matrix has been substantially enriched in silica as well as incompatible elements (Bebout and Barton, 1993). Even volumetrically minor metachert present within the Catalina Amphibolite Unit reveals metasomatism at the peak of subduction metamorphism (Page et al., 2019).
Finally, it is important to emphasize that the metasomatic fluid source that infiltrated the Amphibolite Unit could not have been derived from devolatilization of the presently exposed Catalina Blueschist Unit since the later was accreted ~15–20 m.y. subsequent to peak-grade metamorphism and metasomatism of the Amphibolite Unit (Grove et al., 2008). Rather the low-temperature fluids that infiltrated the Amphibolite Unit were most likely devolatilized from equivalent blueschist facies metasediments that were subducted at deeper levels within the accretionary complex and are no longer preserved.
Structural Architecture of the Catalina Schist
The dominant large-scale structures of the Catalina Schist are late to post-metamorphic thrust faults that abruptly juxtapose units of different metamorphic grade (Fig. 1.3). For example, the basal thrust of the Amphibolite Unit (known as the Ollas thrust) occupies a zone up to a few meters thick of metasomatic rock rich in chlorite, phlogopite, or talc, with rare tectonic blocks. Similarly, the contact that places the Greenschist Unit slices above the Blueschist Unit is a narrow zone of tectonic mélange that juxtaposes rocks with quite distinct mineral assemblages against each other.
The Greenschist-Blueschist thrust and the Ollas thrust both dip regionally to the NE and are folded on a km scale. These large-scale geometrical relationships are compatible with continued deformation within an accretionary complex. In contrast, tectonic boundaries on the NE side of the island are steep or vertical, and the foliation or layering in the metamorphic rocks also dips steeply SW or NE (Fig. 1.3). On the map (Fig. 1.3), these boundaries are shown as thrusts, implying that they have been tilted into a steep orientation. Alternatively, they may be Cenozoic normal faults, dipping SW. While they do not appear to offset the Miocene volcanic rocks, silver mineralization, barite veins, and other evidence of hydrothermal fluid circulation along the faults suggest that they are related to the voluminous Neogene volcanism on Catalina Island.
The Catalina Schist as a whole is overlain unconformably by early Miocene volcanic rocks and some sediments. The eastern portion of the island is pervasively invaded by dikes and hypabyssal intrusions of early Miocene age. Dacite dikes occur sporadically elsewhere on the island. Quaternary surficial deposits consist of extensive landslide deposits, particularly along the steep coastal slopes. On the south side of the island, there is a large body of altered gabbro, which may represent the part of the Coast Range ophiolite; so far this has not been dated. The contacts of the gabbro with the Catalina Schist are not exposed.
Early Exhumation of the Catalina Schist
The greenschist facies overprint and the lack of a blueschist facies overprint in the Amphibolite Unit, together with the predominantly greenschist facies character of the mélange zone along the Greenschist-Blueschist thrust, suggest that the final assembly of the different components of the Catalina Schist accompanied or postdated much of the exhumation. Interpretation of the mechanism of exhumation depends on the overall interpretation of the tectonic setting for the metamorphism of the various components of the terrane, and is discussed in more detail in the section on “Tectonic Model for the Formation of the Catalina Schist,” but the cooling history indicates that exhumation to within ~10 km of the surface was largely complete by the end of the Cretaceous.
Neogene Development of the Continental Borderland and the Final Exhumation of the Schist
Initial contact between the East Pacific Rise and the subducting margin in southern California occurred at ca. 28 Ma near what is now the border between the United States and Mexico (Atwater and Stock, 1998). This contact ushered in a new tectonic regime characterized by crustal extension, right-lateral dextral shear, the formation of slab-free asthenospheric windows, and concomitant volcanism (Atwater and Stock, 1998; Wilson et al., 2005). Beginning around 17 Ma, the southern California borderland experienced a dramatic microplate capture event (e.g., Nicholson et al., 1994) triggered by Pacific–North American shearing along the margin. Late Cretaceous forearc strata and underlying Jurassic and earliest Cretaceous basement rocks were ripped from the western margin of the Peninsular Ranges batholith, rotated clockwise, and translated northwestward (Wright, 1991; Crouch and Suppe, 1993; Bohannon and Geist, 1998; Ingersoll and Rumelhart, 1999), and extension in the wake of this rotating domain may have contributed to the final exhumation of the Catalina Schist. This deformation was further modified by dextral shearing and oblique convergence during the late Miocene, Pliocene, and Quaternary, resulting in a broad zone of deep basins, submerged ridges, and islands referred to as the California Continental Borderland (Howell and Vedder, 1981; Legg, 1991). The Inner Borderland corresponds to the ~100-km-wide part of the offshore region closest to the mainland and is mainly composed of Catalina Schist overlain by Miocene and younger sedimentary and volcanic rocks. The Outer Borderland and western Transverse Ranges include the upper plate rocks that were rifted and rotated away from the California margin during the microplate capture event (Nicholson et al., 1994).
Thermochronologic data from the Catalina Schist exposed on Catalina Island indicate that all map units cooled slowly at > ~10 km depths until ~20 m.y. before present (Grove and Bebout, 1995). As microplate capture and rifting occurred, Catalina Schist basement was tectonically denuded and exhumed within the inner region of the borderland adjacent to the Peninsular Ranges batholith (Woodford, 1925; Stuart, 1979). The coupled microplate capture and resulting extensional event created a unique opportunity to view the accretionary complex that had been previously buried beneath the western margin of the Peninsular Ranges batholith. A somewhat similar Miocene exhumation history exposed subduction-related rocks of the Western Baja terrane in the Vizcaino region of Baja California (Suppe and Armstrong, 1972; Moore, 1986; Sedlock, 1988a, 1988b; Baldwin and Harrison, 1989, 1992).
San Onofre Breccia
Tectonic denudation of the schist exhumed within the inner California borderland produced major syn-orogenic deposits referred to as the San Onofre Breccia (Stuart, 1979; Crouch and Suppe, 1993). Based upon the K-Ar ages of coeval volcanic rocks and biostratigraphy, the depositional age of these deposits ranges from ca. 17 Ma to 10 Ma (Stuart, 1979; Crouch and Suppe, 1993; Bohannon and Geist, 1998). The time-dependent clast composition within the San Onofre Breccia allows a determination of the extent to which the Santa Catalina Island exposures of Catalina Schist represent the much larger area of high P-T rocks that underlie the Inner Continental Borderland of southern California. Clast counts performed by Stuart (1979) appear to be broadly consistent with an unroofing sequence in which amphibolite and epidote-amphibolite grade Catalina Schist and mafic basement with low P-T mineral assemblages (e.g., saussurite gabbro) from the hanging wall were much more abundant at lower stratigraphic levels within the San Onofre Breccia (Stuart, 1979). In contrast, the stratigraphically younger San Onofre Breccia, including deposits laid down atop extended borderland crust underlain by the Catalina Schist, tends to be dominated by lawsonite-blueschist and lower-grade detritus. Based upon these relationships, it is reasonable to conclude that the stacked sequence of high- to low-grade Catalina Schist present on Santa Catalina Island is broadly representative of the Catalina Schist terrane throughout the inner California borderland region.
Miocene Volcanic Rocks
Volcanic rocks in the eastern Santa Monica Mountains dated at 17–15 Ma were an expression of deep crustal magmatism accompanying the earliest extensional tectonism associated with rifting in the Los Angeles basin (Weigand and Savage, 1993; McCulloh et al., 2002). Volcanic rocks are intercalated with San Onofre Breccia and underlie a large extent of the southern California borderland and adjacent Los Angeles basin (Howell and Vedder, 1981; Vedder, 1987; Wright, 1991; Crouch and Suppe, 1993; Bohannon and Geist, 1998; ten Brink et al., 2000; McCulloh et al., 2002). A number of enigmatic crater structures mapped throughout the inner southern California borderland may indicate that widespread explosive volcanism and caldera collapse occurred throughout the highly extended terrane (e.g., Legg et al., 2004).
A major Miocene volcanic center underlies the southeastern portion of Catalina Island and illustrates well the stratigraphic relationships found throughout the inner southern California borderland. A large, hypabyssal complex referred to as the Catalina pluton consists of fine-grained quartz diorite to granodiorite and has been dated by K-Ar methods at 19 Ma (Forman, 1970). The hypabyssal complex contains xenoliths of Catalina Schist, and is pervasively invaded by vertical sheeted sequences of andesitic to dacitic dikes that are locally intercalated with San Onofre Breccia (Boundy-Sanders et al., 1987). On the east end of the island (Seal Rocks), an active quarry operation reveals that San Onofre Breccia, deposited atop a marine bench incised into the hypabyssal pluton, is in turn intruded by dikes (Vedder et al., 1979). Elsewhere on the island, dacitic rocks dated at 15–14 Ma by K-Ar methods form domelike masses that are flanked and overlain by 15–12 Ma andesitic flows, some of which may have been subaerial (Vedder et al., 1979; Boundy-Sanders et al., 1987). Near the isthmus, small lenses of San Onofre Breccia, coarse-grained volcaniclastic strata, and middle Miocene diatomaceous shale and lapilli tuff drape over volcanic flows. Collectively, these observations indicate deposition within an active extensional environment (Vedder et al., 1979; Boundy-Sanders et al., 1987).
Present Tectonic Setting
Over the past 20 m.y., the diffuse plate boundary through coastal southern California has evolved from accommodating microplate capture to a complex dextral transform boundary (Bohannon and Geist, 1998). Ingersoll and Rumelhart (1999) have proposed a three-stage evolutionary history for the Los Angeles basin. During the initial phase (18–12 Ma), the western Transverse Ranges underwent transrotation during the microplate capture event (Nicholson et al., 1994). The second phase persisted from 12 to 6 Ma and was characterized by transtension. The timing of this phase corresponds to slip along the San Gabriel fault (Crowell, 1975) and major sedimentation within the Los Angeles basin (e.g., Wright, 1991). The third stage occurred after significant slip along the southern branch of the modern San Andreas fault began at ca. 6 Ma (Dickinson, 1996).
North-northwest–striking slip faults oriented subparallel to the San Andreas fault currently define the structural grain of the borderland region and host much of the seismic activity (Sorlien et al., 2015). The most prominent of these, the dextral-slip, NNW-trending Newport–Inglewood–Rose Canyon fault, runs along the southern California coast and has a slip rate of 1.07 ± 0.03 mm/yr (Rockwell et al., 1992). Other significant, similarly oriented faults situated progressively further offshore include the Palos Verdes fault, San Pedro basin fault, and the San Clemente Island fault (Fisher et al., 2004). The high-angle, dextral, NNW-trending faults interact in complex ways with the middle Miocene detachment faults. For, example, both the Oceanside thrust and a similar structure further offshore (Thirty Mile Bank thrust) may have originated as normal faults and were then reactivated as blind thrusts (Rivero et al., 2000). While considerable geodetic shortening has been proposed for these structures, their reactivation as blind thrusts remains controversial (Sorlien et al., 2015).
A series of transpressional basement uplifts has been produced in the coastal southern California region as a result of the late Pliocene–Quaternary “Pasadena orogeny.” These include the San Gabriel Mountains (e.g., Blythe et al., 2000) and less significant, localized features such as Palos Verdes Peninsula (Woodring et al., 1945), Catalina Island (Castillo et al., 2019), and San Clemente Island (Ward and Valensise, 1996). On Palos Verdes Peninsula and San Clemente Island, well-developed marine terraces document recent uplift. In the case of Palos Verdes Peninsula, deformation related to the Palos Verdes reverse fault has produced 13 individual marine terraces, dating back to ca. 2.5 Ma, that encircle the peninsula in a bathtub ring configuration (Ward and Valensise, 1994). Analysis of the elevations of terrace remnants indicates a displacement model involving 3.0–3.7 mm yr–1 of oblique, dextral/reverse slip on a fault dipping 67° at 6–12 km depth beneath the peninsula (Ward and Valensise, 1994). Similarly, on San Clemente Island, 3 km of compression normal to the N150E fault strike over the past 2–5 m.y. has produced 1700 m domal uplift of the San Clemente anticlinorium (Ward and Valensise, 1996). Analysis of uplifted marine terraces on San Clemente Island indicates current uplift rates of 0.2 and 0.5 mm yr–1. In contrast, Catalina Island currently appears to be subsiding (Castillo et al., 2019). Submerged paleoshorelines and terraces surrounding Santa Catalina Island and the Pilgrim/Kidney Banks in the California Continental Borderland demonstrate significant late Quaternary tectonic subsidence. Based upon a host of remotely operated vehicle subsurface mapping and sample recovery from these shorelines, Catalina Island is indicated to have been tilting north and subsiding together with its surrounding platform at 0.08–0.27 mm/yr over the past 1.15 Ma (Castillo et al., 2019). Collectively, these observations indicate that a dynamic topographic environment persists within the Inner Borderland region.
DESCRIPTION OF FIELD-TRIP STOPS
Please note that rock collecting is not allowed on the island without a research permit from the Catalina Conservancy.
Stop 1. Catalina Harbor, South Side (33°25′40.47″N, 118°30′22.29″W) (see Fig. 1.1 for location)
Metagraywackes of the Catalina Blueschist Unit. This is the lowest tectonic unit on Catalina Island, and in this area the rocks carry the relatively low-grade assemblage quartz + white mica + chlorite + lawsonite + albite. The disruptive structure and abundant quartz veining are characteristic of the deformational style in much of this unit. Thick-bedded, coarse-grained metasandstone has a detrital texture visible in outcrop, and there are some thin pebbly bands. Finer-grained rocks have a slaty-type cleavage, produced by pressure solution, and bedding is largely obscured by boudinage and quartz veining. The veins indicate high fluid pressure during deformation, which may reflect compaction, dehydration reactions such as clays → white mica, or fluids moving up the subduction channel from greater depths.
Stop 2. View Stop, Junction of Middle Canyon Road and Escondido Ranch Road (33°22′46.60″N, 118°28′39.99″W) (see Fig. 1.3 for location)
This stop provides a good overview of the metamorphic units in the center of the island. To the north is Little Harbor, where we will spend part of the morning looking at rocks of the Catalina Blueschist Unit, and beyond that Little Springs Canyon, where we will see rocks of the overlying Catalina Greenschist Unit on Day 3. To the northwest, the hills are formed from resistant rocks of the Amphibolite Unit; further to the right Catalina Airport, which is built on the ultramafic mélange that forms the structurally highest part of the Catalina Schist, is visible. We will spend time in that area this afternoon. To the west the peaks of Black Jack (with the communications tower) and Orizaba are made up of early Miocene volcanic rocks. We are currently standing on one of the dissected geomorphic surfaces, dipping gently SW, which characterize the landscape in this part of the island.
Stop 3. Little Harbor (33°23′13.54″N, 118°28′22.98″W) (see Fig. 1.3 for location)
Walk to the north end of the beach to see a metaconglomerate body that forms part of the Blueschist Unit. Rounded clasts are formed mainly from plutonic and volcanic rocks, likely sourced from the magmatic arc to the east. Clasts and matrix are metamorphosed in the glaucophane-lawsonite schist facies: hornblende in metagabbro and pyroxene in metadacite are replaced by glaucophane; plagioclase is replaced by lawsonite. Bedding is difficult to see, but is vertical. The conglomerate shows little penetrative deformation. If the tide allows, we can scramble around to the west side of the conglomerate body, and see its contact relations with deformed metagraywacke sandstone and shale. The conglomerate clearly formed a resistant body immersed in a viscous matrix, which has been invoked as a possible explanation for seismic tremor and slow slip.
Walk south to the promontory in the center of the harbor. On the beach is an exposure of coarse-grained mica schist belonging to the Catalina Greenschist Unit. Note the marked contrast in metamorphic grain size and degree of recrystallization in comparison to the metasediments in the Blueschist Unit. This outcrop forms part of a small klippe of the Greenschist Unit extending inland from the harbor. The tectonic contact between the two units is visible just to the south. The promontory is made up of mafic blueschist interlayered with metasediments of the Blueschist Unit, and must directly underlie the thrust. The mafic blueschist contains the assemblage glaucophane + lawsonite + sphene (Fig. 1.6B), characteristic of this unit.
Stop 4. Little Harbor, South Side (33°22′51.82″N, 118°28′33.59″W)
From the parking lot at the south end of the beach, we will hike along a scrambly trail that runs ~250 m SW along the coast. Take care: the trail has been washed out in places. The route takes us through exposures of mafic blueschist, metasedimentary rocks (primarily dark graphitic phyllite), and a variety of fragmental rocks that have been described as mélange. In the mélange, note flattened fragments of metasandstone, black phyllite, mafic blueschist, and various soft pale greenish rocks that have a broadly ultramafic composition (Fig. 1.5C). The latter are composed of varying proportions of Mg-rich chlorite, Mg-glaucophane, serpentine, and talc, and clearly reflect substantial metasomatic alteration of the original rock. Small flecks of bright green fuchsite (Cr-mica) or mariposite (Cr-chlorite) may have formed from primary chrome spinel. The fragments and matrix of the mélange have all been metamorphosed in the blueschist facies. We will discuss possible origins for this type of mélange.
In a cove at the end of the trail (33°22′51.82″N, 118°28′33.59″W), we will see metabasalt with pillow texture, nodular metachert, and fallen blocks of metaconglomerate showing varying degrees of deformation (Fig. 1.5B). There is also an excellent exposure of a sheeted vein complex in graphitic phyllite (Fig. 1.5D). This may mark a fault or shear zone within the unit, formed during subduction and accretion.
Stop 5. Ben Weston Beach (33°22′16.70″N, 118°28′49.13″W) (see Fig. 1.1 for location)
This requires a fairly steep descent from the parking area to the beach. On the way down, note exposures of various rocks attributable to the Blueschist Unit, including metagraywacke, black phyllite, glaucophane schist, serpentinite, and talc or chlorite-rich schist. The beach is shown on topographic maps as Mills Landing, but is known to islanders as Ben Weston Beach. On the north side, there is an impressive cliff exposure of metasedimentary rocks showing a large-scale overturned fold in metagraywacke sandstone (greenish-buff color) and shale (dark phyllites). Note the concentration of quartz veins in the phyllites, reflecting their differing mechanical properties. The fold is associated with the primary schistosity in the rocks; this has been modified by small-scale folding and crenulation, producing interference structures between the two phases of folding. Unfortunately, accessible exposures of these small-scale structures were destroyed in a landslide.
On the south side of the beach, small-scale second-generation folds are developed, folding the primary schistosity. The metasediments here are intensively veined: vein minerals include quartz, albite, and carbonate minerals. Some of the quartz veins reflect hydraulic fracture, others appear to be replacive, to the point that original sandstone beds appear to have been largely replaced by quartz. Albite veins are largely replacive; they are commonly black, due to inclusions of graphite, and the graphite inclusions delineate ghost schistosity and crenulation cleavage inherited from the original rock.
We will follow a segment of the Catalina Trail down the hill to the south for 200 m, and see a steatite bowl factory set up by Native Americans prior to depopulation of the island in the nineteenth century. Steatite figurines and bowls (known by the Spanish term ollas) were traded extensively around southern California. The steatite is massive talc rock forming part of the ultramafic mélange at the top of the Catalina Amphibolite Unit. We will then walk further down the trail, passing a number of garnet hornblendite blocks, and find an exposure in a dry gully of the deformed mélange matrix, composed of various proportions of chlorite, serpentine, talc, and ortho- and clinoamphiboles. The origins of the mélange remain a puzzle, and most of the rocks have compositions suggesting extensive metasomatism, but it is generally interpreted as being a result of tectonic mixing in a subduction channel. We will also see blocks of massive serpentinized peridotite. It is plausible that the peridotite originally formed part of the subduction zone hanging wall (i.e., the mantle wedge). P-T conditions obtained from both the garnet hornblendite blocks and the mélange matrix are in the range 8–11 kbar and 640–750 °C.
Turn left onto Empire Landing Road (graded dirt) from Airport Road (paved).
Stop 7. Roadcut on North Side of Catalina Airport (33°24′28.48″N, 118°24′41.08″W)
This roadcut shows folded metachert of the Catalina Greenschist Unit. We have now crossed one of the steep NW-trending faults (the Airport fault) along the northern side of the island, possibly of Tertiary age, which juxtaposes rocks of the Greenschist Unit against the ultramafic mélange (Fig. 1.3). The metachert unit here is tightly folded, and in thin-section, shows a mylonitic microstructure produced by dynamic recrystallization accompanying dislocation creep in quartz (Fig. 1.6F). Metachert in Cottonwood Canyon, just beneath the Ollas thrust at the base of the Amphibolite Unit, shows similar microstructures, which may be related to the thrust. The metachert commonly carries sodic amphibole, garnet, and piemontite. Pelitic schist nearby has garnet and biotite. This stop provides excellent views down over the Blueschist and Greenschist Units on the northern slopes of the island, and across the San Pedro Channel to the mainland. On a clear day, it is possible to see the San Gabriel, San Bernadino, and Santa Ana Mountains, and even San Jacinto Mountain, 130 miles to the east.
Stop 8. Contact between Ultramafic Mélange and Coherent Mafic Amphibolite along Sheep Chute Jeep Trail (33°24′29.11″N, 118°25′54.13″W) (see Fig. 1.3 for location)
If you are visiting this stop without the 2020 GSA Cordilleran Section Meeting leaders, please obtain a Catalina Island Conservancy permit (https://catalinaconservancy.org/community/) before hiking in the area.
We will drive 1.54 miles (2.47 km) through the amphibolite facies mélange to the intersection of Empire Landing Road and Sheep Chute jeep trail (33°24′29.11″N, 118°25′54.13″W). Visible blocks include garnet amphibolite and silicified ultramafic rocks set in chlorite-rich and talc-rich mélange respectively. After parking, we will walk an additional 250 m along Sheep Chute jeep trail through the amphibolite facies mélange. In addition to the previously mentioned lithologies, blocks of garnet-bearing metachert(?) will be encountered. The contact between the amphibolite facies mélange and structurally underlying coherent amphibolite facies rocks is poorly exposed. The structurally highest coherent unit is a subhorizontal slice of quartz – oligoclase – biotite ± garnet ± kyanite – bearing metasediment that is locally cut by 2–20-cm-thick quartz-oligoclase dikes. These metasediments are partially retrogressed to greenschist facies and are typically highly weathered. Walk another ~300 m along the Sheep Chute jeep trail to encounter strongly foliated metagabbro beginning ~550 m west of where we parked. The metagabbro dips shallowly and underlies the area south of the airport. Spend some time observing these relationships and return to the vehicle.
Stop 9. Little Springs Canyon (33°24′31.15″N, 118°28′6.42″W) (see Fig. 1.3 for location)
We will park in a pull-out on the road from Two Harbors to Little Harbor. From here we have a good view east of the NW end of the ridge, which is held up by the Catalina Amphibolite Unit. The two hills directly to the east are made up of massive serpentinized peridotite, overlying mafic amphibolite. This in turn rests directly on rocks of the Blueschist Unit along the Ollas thrust. The Ollas thrust here is folded into a km-scale synform, plunging gently south. The synform is asymmetric, with a gently dipping west limb and a steep east limb.
We will scramble down the hillside through the cactus into Little Springs Canyon, following the thrust contact between the Greenschist Unit and the underlying Blueschist Unit, which is marked here by a zone of various metasomatically altered ultramafic rocks with tectonic blocks of amphibolite facies rocks. In the creek bed, we’ll see a large tectonic block of garnet hornblendite with eclogitic layers—watch out for petroglyphs! (Reminder: No rock collecting is permitted on the island.) This hornblendite rests on crenulated phyllite of the Blueschist Unit, which gave a laser Raman temperature on carbonaceous material of 337 °C.
We will then make our way ~300 m south down the canyon, which requires some care, as the sides are steep, there is a lot of brush, and there are several deep pools filled with mud or stagnant water. We pass south, downstream but structurally upward, into a klippe composed of albite-epidote amphibolite facies rocks, which forms one of the thrust slices collectively known as the Catalina Greenschist Unit. The internal structure of the klippe is shown on the detailed map and cross section in this guide (Figs. 1.8 and 1.9). We pass through metasedimentary mica schist, which gave a laser Raman temperature estimate of 573 °C, and then into metamafic rocks, with pale-green hornblende, epidote, albite/oligoclase, and sphene. The epidote concentrations in places appear to outline relict pillow structure.
Structurally beneath the metamafic rocks, there is a beautifully exposed complex antiform in layered metachert, with parasitic folds and a distinct axial-plane schistosity. The antiform plunges strongly west, and the metamafic rocks structurally overlie the cherts, so the antiform may be downward facing. The metachert locally contains one or more of sodic amphibole, biotite, garnet, and stilpnomelane, but these minerals are commonly retrogressed to chlorite.
Eighty meters further south, downstream, we go back through the metamafic rocks, and then into another antiform of metachert. This has been quite strongly refolded, so that most of the visible folds are second generation N-vergent folds with gently S-dipping axial planes. It is possible to find refolded first-generation folds, and nice examples of the first schistosity folded by the second-generation folds. The metachert layer can be mapped over the whole area of the klippe, and exhibits numerous folds of both generations (Fig. 1.9). Note that although the structure is quite complicated, it does not have the disruptive character exhibited by the Blueschist Unit, and the deformation is entirely ductile.
Return to vehicle and drive to next stop.
Stop 10. “Little Canyon” (33°23′22.47″N, 118°28′20.46″W) (see Fig. 1.3 for location)
From the parking area, we walk ~400 m NE up a small dry canyon north of Little Harbor. This has very nice exposure of mafic blueschist of the Catalina Blueschist Unit with some metasedimentary interlayers. The primary schistosity has been intensely microfolded or crenulated (Fig. 1.6C). Refolded first-generation folds are locally visible, and both small- and large-scale second-generation folds are associated with the crenulations. The mafic blueschist carries glaucophane, lawsonite, and sphene, and the foliation defined by these minerals is kinked and folded by the second-generation folds. Schistosity surfaces commonly carry a mineral lineation and one or two crenulation lineations. The outcrop is fairly continuous for several hundred meters, illustrating a coherent style of deformation that contrasts with the disruptive structure exhibited by the rocks in the Little Harbor area.
Stop 11. Big Fisherman Cove (33°26′42.39″N, 118°29′0.24″W) (see Fig. 1.1 for location)
Depending on time, we will examine an exposure of early Miocene volcaniclastic rocks on the north side of Big Fisherman Cove.
TECTONIC MODEL FOR THE FORMATION OF THE CATALINA SCHIST
Platt (1975), Peacock (1987), and others originally interpreted the Catalina Schist as an inverted metamorphic sequence formed during nascent subduction and accretion beneath unrefrigerated mantle lithosphere in a progressively cooling subduction zone, and applied this interpretation to the metamorphism of high-grade tectonic blocks within the Franciscan Complex in the California Coast Ranges. The application of the subduction initiation concept to the Franciscan Complex has largely stood the test of time. Early Jurassic isotopic ages from garnet amphibolites and eclogites within the Franciscan Complex can readily be reconciled with a subduction initiation event coeval with the formation of the Early–Middle Jurassic Coast Ranges ophiolite at around 165 Ma (Cloos, 1985; Anczkiewicz et al., 2004; Mulcahy et al., 2018).
While a nascent subduction model is appealing for the high-grade blocks in the Franciscan Complex, a subduction initiation event at ca. 115 Ma represented by high-grade rocks of the Catalina Schist is much more difficult to reconcile with the Jurassic and Cretaceous emplacement history of the adjacent Peninsular Ranges batholith (Grove and Bebout, 1995). The mid-Cretaceous plate kinematics involving the Farallon, Kula, and Pacific plates are not definitive (Engebretsen et al., 1986; Stock and Molnar, 1988; Müller et al., 2008). Nevertheless, all available studies indicate that continuous subduction of comparatively old oceanic crust likely prevailed along the southwestern North American margin throughout the mid-Cretaceous. While the thermal effect of ridge subduction involving the Kula and Farallon plates to form the high-grade rocks of the Catalina Schist is conceivable (e.g., Engebretsen et al., 1986; Stock and Molnar, 1988), little supporting evidence exists for such a scenario.
Important clues that the high-grade and low-grade rocks of the Catalina Schist may be genetically unrelated and formed in distinct environments include: (1) the lack of blueschist facies overprinting of the high-grade rocks; and (2) the 15–20 m.y. gap in time between amphibolite facies and blueschist facies metamorphism. Widespread blueschist facies overprinting of amphibolite and epidote amphibolite facies rocks occurs within high-pressure/low-temperature accretionary complexes such as the Franciscan Complex (Cloos, 1985; Moore and Blake, 1989; Wakabayashi, 1990). Simple thermal models of early-stage subduction that are capable of producing inverted metamorphism that varies from amphibolite facies to blueschist facies at 1.0 GPa (~40 km) depths require substantial subduction refrigeration over very short time scales and result in counter-clockwise pressure-temperature paths (Peacock, 1987). In the case of the Catalina Schist, geothermal gradients would have to decrease by more than a factor of two (from ~18–8 °C/km) over a 15–20 m.y. time scale. Under such conditions, it is very difficult to avoid blueschist facies overprinting if all rocks remained at ~30 km depths (e.g., Ernst, 1988; Wakabayashi, 1990; Plunder et al., 2016; Cordova et al., 2018).
Modern computation efforts that take mantle corner flow and temperature-dependent rock rheology into account produce highly nonlinear geotherms along the top of the subducting slab (Syracuse et al., 2010), and tend to yield clockwise curvature of computed P-T paths. While this permits production of blueschist facies through amphibolite facies conditions at the same approximate level, the rocks exhibiting these varying grades of metamorphism are spatially well separated. Moreover, the depth at which geotherms become nonlinear is typically 2.5–2.0 GPa (~70–88 km) or twice as deep as the pressure estimates for the Catalina Schist.
In a systematic global study of subduction zone metamorphism, however, Penniston-Dorland et al. (2015) concluded that P-T conditions recorded by exhumed subduction accretionary complexes are generally warmer than those recorded by thermokinetic models. This discrepancy may reflect petrologic overestimation of temperatures due to thermal overprinting during exhumation, and/or failure of numerical models to appropriately deal with shear heating, hydration reactions, and fluid and rock advection (Penniston-Dorland et al., 2015; Kohn et al., 2017). Despite all of this, the Amphibolite Unit of the Catalina Schist experienced metamorphic temperatures that are exceptionally hot for accretionary complex rocks worldwide (e.g., Penniston-Dorland et al., 2015).
Given the available constraints, it is evident that the amphibolite facies and blueschist facies rocks of the Catalina Schist could not have formed at the same time in a single tectonic environment. Grove and Bebout (1995) and Grove et al., (2008) favored a model in which the high- and low-temperature rocks of the Catalina Schist formed in different locations and were later juxtaposed. In this model, amphibolite facies and epidote amphibolite facies rocks represent forearc basement and sedimentary cover that was subducted beneath mantle lithosphere along the leading edge of the Peninsular Ranges batholith between 115 and 100 Ma and accreted at ~40 depth. This event collapsed the Early Cretaceous forearc region within southern and northernmost Baja California by up to ~50 km, and caused a major eastward relocation of the locus of arc magmatism after 100 Ma by an equivalent distance. This shortening event resulted in erosional denudation of the Peninsular Ranges batholith that flooded the forearc, and ultimately led to accretion of the epidote-blueschist facies, lawsonite-blueschist facies, and lower-grade units of the Catalina Schist between ca. 100–95 Ma. Below, we describe Grove et al.’s (2008) model in greater detail.
Continuity of Jurassic–Early Cretaceous Ophiolite Basement along the Cordilleran Margin
Ophiolite basement of predominantly Middle Jurassic age crops out discontinuously along a 2300-km-long segment of the Cordilleran margin from southern Oregon to the Magdalena Bay region of Baja California (Suppe and Armstrong, 1972; Jones et al., 1976; Kilmer, 1979; Kienast and Rangin, 1982; Rangin et al., 1983; Kimbrough, 1985, 1989). On the Vizcaino peninsula and nearby Cedros Island of west-central Baja California, both Late Triassic and Middle Jurassic ophiolite basement underlie the forearc region. These excellent exposures provide an extraordinarily clear record of supra-subduction zone ophiolite generation (Pessagno et al., 1979; Moore, 1985; Kimbrough, 1985; Kimbrough and Moore, 2003; Dektar and Chapman, 2018) (Fig. 2.1). The Middle Jurassic ophiolite basement on Cedros Island yields 173 ± 2 Ma U-Pb zircon ages (Kimbrough and Moore, 2003). In both the Vizcaino and Cedros Islands areas, this supra-subduction basement structurally overlies blueschist-bearing subduction assemblages (Suppe and Armstrong, 1972; Moore, 1986; Sedlock, 1988a, 1988b; Baldwin and Harrison, 1989, 1992). Ophiolite basement in west-central Baja California is depositionally overlain by a thick sequence of Jurassic–Early Cretaceous volcanic and clastic sediments (Busby-Spera, 1988; Critelli et al., 2002; Busby, 2004).
In northern California and southern Oregon, a similar relationship exists between the Coast Ranges ophiolite and the subduction-related Franciscan Complex (Bailey et al., 1970; Ernst, 1970; Suppe and Armstrong, 1972; Jones et al., 1976; Shervais et al., 2004; Hopson et al., 2008). The age of the Cedros ophiolite overlaps the range of crystallization ages determined for both the Coast Ranges ophiolite of central and northern California (172–160 Ma; Shervais et al., 2004; Hopson et al., 2008). Similarly, Late Jurassic–Early Cretaceous basal Great Valley Group was deposited upon the Coast Ranges ophiolite (Ingersoll, 1979; DeGraaff-Surpless et al., 2002; Orme and Surpless, 2019).
Underthrusting of the Jurassic–Early Cretaceous Forearc beneath the Peninsular Ranges Batholith (34–32°N)
The continuity of geophysical trends along the Baja California margin allows the basement geology of the Vizcaino region to be extrapolated northward (Langenheim et al., 2014). Similarly, the linear trend of Coast Range ophiolite outcrops projects southward toward southern California. It is therefore reasonable to presume that in Early Cretaceous time, a continuous belt of Jurassic to Early Cretaceous ophiolite basement structurally underlain by a subduction complex, and depositionally overlain by a Late Jurassic–Early Cretaceous forearc basin, existed within the southern California region (e.g., Jones et al., 1976).
Mafic basement assemblages are present in the southern California region on Santa Cruz Island (Weaver and Nolf, 1969; Hill, 1976; Mattinson and Hill, 1976). The Willows Plutonic Complex on Santa Cruz Island consists mostly of hornblende diorite and quartz diorite and less abundant gabbro, together with minor plagiogranite and ultramafic rocks (Hill, 1976). Compositionally similar, albeit highly altered rocks are present on Catalina Island (Hill, 1976), as basement encountered in drillholes into the southwestern Los Angeles basin (Schoellhamer and Woodford, 1951; Yeats, 1973; Sorenson, 1985, 1988b), and as clasts within the San Onofre Breccia (Stuart, 1979). Within the outer California Continental Borderland, offshore drilling and seismic exploration detect mafic basement beneath forearc strata (Bohannon and Geist, 1998; ten Brink et al., 2000). A U-Pb zircon age of 162 Ma was reported by Mattinson and Hill (1976) from plagiogranite in the Willows Complex. Here we report three new laser ablation zircon U-Pb ages from gabbro samples in the Willow Complex that yield indistinguishable ages of ca. 168 ± 3 Ma (Table 1). These results support correlation of the mafic basement rocks within the southern California borderland to the Coast Range Ophiolite and similar basement underlying west-central Baja California. Finally, while Late Jurassic and Early Cretaceous volcanic rocks and forearc strata are uncommon throughout the southern California borderland, sedimentary and volcanic units of this age do occur as metamorphosed wall rocks. They include the Upper Jurassic Peñasquitos Formation (Santa Ana Mountains), Upper Jurassic Santa Monica Formation (Santa Monica Mountains), and poorly dated, volcanogenic schist on Santa Cruz Island (Hill, 1976; Kimbrough et al., 2014).
The relationships described above led Grove et al. (2008) to represent the early Cretaceous margin of the northern Peninsular Ranges batholith as depicted in Figure 2.2A. In Grove et al.’s (2008) model, the Late Jurassic–Early Cretaceous forearc region of southern California and northernmost Baja California began to be underthrust to ~40 km depth beneath the leading edge of the Peninsular Ranges batholith after 120 Ma (Fig. 2.2B). Grove et al. (2008) associated the Ultramafic mélange within the Amphibolite Unit with lithospheric mantle within a shear zone developed beneath the western batholith. They furthermore associated the gabbroic rocks of the Catalina Schist amphibolite unit to the Coast Range ophiolite, and the structurally overlying metasedimentary rocks to Late Jurassic–Early Cretaceous forearc strata. Underthrusting of the forearc basement to a stalled position beneath the western margin of the magmatic arc explains the heat source required to produce amphibolite facies metamorphism and anatexis within the Catalina amphibolite unit (e.g., Sorensen and Barton, 1987) (Fig. 2.2B). It also accounts for why the amphibolite unit experienced protracted greenschist facies metamorphic conditions between 115 and 105 Ma (Fig. 2.2C). Finally, Grove et al. (2008) pointed out the cratonal nature of the detrital zircon provenance signature of the amphibolite facies and epidote amphibolite facies metasediments from the Catalina Schist. This provenance signature is similar to that of the Upper Jurassic Peñasquitos and Santa Monica Formations (e.g., Kimbrough et al., 2014).
The timing of shortening and underthrusting indicated in Fig. 2.2B is supported by considerable evidence for shortening within the northern Peninsular Ranges batholith between ca. 120–105 Ma (Gastil et al., 1981; Todd and Shaw, 1985; Silver and Chappell, 1988; Thomson and Girty, 1994; Schmidt and Paterson, 2002; Todd et al., 2003; Morton et al., 2014). The Cuyamaca-Laguna Mountain shear zone of Todd and Shaw (1985) and Thomson and Girty (1994) is one such prominent fault within the northern Peninsular Ranges (Fig. 2.2B). Schmidt and Paterson (2002) have described a similar high-strain axial fan structure further south. Mid-Cretaceous shortening and underthrusting of the forearc could have been triggered by lateral expansion related to batholith emplacement, coupled with collapse of thin, hot backarc crust that had separated the Peninsular Ranges batholith from the southwestern craton margin in the Early Cretaceous (Busby, 2004).
Based upon detrital zircon results, epidote-amphibolite rocks included within the Greenschist Unit were accreted after 113 ± 3 Ma or several million years after peak grade recrystallization in the overlying amphibolite unit (Fig. 2.2C). Slow-cooling through the greenschist facies, rather than blueschist facies overprinting, is expected if the amphibolite unit was metamorphosed close to the magmatic arc and far removed from the subduction zone. Greenschist facies metamorphic conditions likely persisted until ca. 97 Ma, based upon muscovite and phengite 40Ar/39Ar results from amphibolite unit metasediments and the epidote-amphibolite metagraywackes of the Greenschist Unit.
Relationship to Intra-Arc Thrusting along the Mátir Thrust Further South (32–29°N)
Early Cretaceous volcanic cover associated with the northern Peninsular Ranges (i.e., “Santiago Peak Volcanics”) depositionally overlies coeval plutons of the western Peninsular Ranges batholith with little deformation evident (Herzig and Kimbrough, 2014). Alternatively, age equivalent volcanic rocks south of the Agua Blanca fault at 32°N are penetratively deformed and underthrust beneath a regionally extensive feature referred to as the Mátir thrust (Fig. 2.1). The Mátir thrust extends south of 32°N along a ~350 km strike length of the Peninsular Ranges batholith (Johnson et al., 1999; Wetmore et al., 2002; Schmidt and Paterson, 2002; Busby et al., 2006; Alsleben et al., 2014; Schmidt et al., 2014; Wetmore et al., 2014). The thrust juxtaposes the 130–100 Ma Alisitos volcanics and associated plutons of the western Peninsular Ranges batholith beneath plutons of the eastern batholith. Structural analysis performed on plutons well-dated by zircon U-Pb methods indicates that the tectonic shortening along the Mátir thrust occurred between 115 and 100 Ma (Alsleben et al., 2014; Schmidt et al., 2014).
The Mátir thrust has previously been considered to represent a fundamental suture between the continent and a volcanic island arc prior to 100 Ma (Johnson et al., 1999; Wetmore et al., 2002; Schmidt and Paterson, 2002; Busby et al., 2006; Alsleben et al., 2014; Schmidt et al., 2014; Wetmore et al., 2014). However, the coeval nature and compositional similarity of the undeformed Santiago Peak to the deformed Alisitos arc segment indicates that both are part of the same volcanic arc. Moreover, Proterozoic zircon within sedimentary rocks intercalated with the volcanic arc rocks (Wetmore et al., 2002), and older basement overlain by it (Kimbrough et al., 2014) indicate that the volcanic arc was never widely separated from the southwestern North American margin.
Wetmore et al. (2014) have pointed out that the Santiago Peak volcanic arc segment of the Peninsular Ranges essentially corresponds to the portion of the batholith that is considered by Grove et al. (2008) to have overthrust its forearc region. Wetmore et al. (2014) proposed that Early Cretaceous shortening and underthrusting of the forearc between 34 and 32°N may have been transferred into the magmatic arc by the ancestral Agua Blanca fault, to be re-expressed as the Mátir thrust south of 32°N. The observation that the Mátir thrust dies out south of 29°N helps explain the full preservation of the Jurassic–Early Cretaceous forearc in the Vizcaino region (Fig. 2.1).
Mid-Cretaceous Eastern Relocation and Composition Transformation of the Peninsular Ranges Batholith
Shortening of the forearc region ultimately resulted in an abrupt eastward relocation of the axis of magmatism by ~50 km beginning ca. 100 Ma. Prior to 100 Ma, the western Peninsular Ranges batholith had been constructed atop relatively primitive mafic crust from a shallowly sourced, plagioclase-bearing source region that produced a broad range of calc-alkaline plutons ranging in composition from gabbro to tonalite to granite (Silver and Chappell, 1988; Todd et al., 2003; Lee et al., 2007; Premo et al., 2014; Kimbrough et al., 2015). After 100 Ma, a new suite of large (1000–3000 km2), compositionally zoned, deeply sourced (e.g., garnet stable) tonalite-trondhjemite-granodiorite plutons, informally referred to as the “La Posta” suite, was intruded ~50 km to the east during a high-flux (e.g., 100 km3/km strike/m.y.) event centered at 95 ± 4 m.y. (Silver and Chappell, 1988; Walawender et al., 1990; Tulloch and Kimbrough, 2003) (Fig. 2.2D).
Emplacement of the voluminous La Posta plutons triggered significant denudation and deep erosion of the eastern batholith that flooded the forearc region with Cenomanian–Turonian detritus, along a 750 km strike length of the batholith from the Santa Ana Mountains to the Vizcaino region (Kimbrough et al., 2001; Grove et al., 2003a; Sharman et al., 2015). Abundant Cenomanian and younger forearc strata onlap the western margin of the northern Peninsular Ranges batholith (Woodring and Popenoe, 1942; Yerkes et al., 1965; Flynn, 1970; Nordstrom, 1970; Peterson and Nordstrom, 1970; Kennedy and Moore, 1971; Sundberg and Cooper, 1978; Schoellhamer et al., 1981; Nilsen and Abbott, 1981; Bottjer et al., 1982; Bottjer and Link, 1984; Fry et al., 1985; Girty, 1987; Bannon et al., 1989). Thick sections of Upper Cretaceous strata also occur throughout the Outer Borderland (Howell and Vedder, 1981; Vedder, 1987; Bohannon and Geist, 1998) (Fig. 2.1). This Cenomanian and younger sediment likely filled the trench, was subducted, and ultimately accreted at depth to form the epidote blueschist, lawsonite blueschist, and lower-grade rocks of the Catalina Schist by 95 Ma (Grove et al., 2008) (Fig. 2.2D). The detrital zircon provenance signature of all lower-grade rocks of the Catalina Schist strongly resembles the age distribution of mid-Cretaceous plutons of the Peninsular Ranges batholith, and lacks the cratonal provenance of higher-temperature metasediments within the Catalina Schist (Grove et al., 2008).
Taylor (1986), Gromet and Silver (1987), and Silver and Chappell (1988) identified the source of the eastern zone La Posta plutons as deeply accreted mafic and sedimentary rocks of a deep-seated accretionary complex, on the basis of distinctive isotopic, trace element, and whole-rock compositional characteristics. Grove et al. (2008) viewed crustal shortening and subduction erosion of the forearc region as the trigger for generating La Posta magmatism. Specifically, eastern relocation of the batholith provided access to warmer mantle via corner flow, just as a high flux of low-density (i.e., quartzofeldspathic) Late Jurassic–Early Cretaceous forearc crust was subducted and supplied to a source region at the base of the crust in a relamination process (e.g., Hacker et al., 2011) (Fig. 2.2D).
Pluton emplacement within the eastern Peninsular Ranges batholith ceased after ca. 88 Ma as arc magmatism migrated eastward into mainland Mexico (Silver and Chappell, 1988; Premo et al., 2014). Extinction of the mid-Cretaceous continental margin magmatic arc occurred as a result of the flattening of the angle of subduction during the Laramide orogeny after ca. 85 Ma (Jacobson et al., 2011; this study) (Fig. 2.2E). Laramide continental margin deformation along the northern extent of the Peninsular Ranges batholith may have involved subduction of the conjugates to the Shatsky and Hess oceanic plateaus (Saleeby, 2003; Liu et al., 2010). We propose that juxtaposition of the higher-grade units of the Catalina Schist with the subduction complex was a Laramide event that occurred at ~10–30 km depths well after accretion and cooling of the Catalina blueschists had occurred. Juxtaposition of the units may have resulted from tectonic denudation via low-angle extensional faulting of overthickened crust (e.g., Platt, 1986) (Fig. 2.2E).
RELATIONSHIP TO THE PELONA, OROCOPIA, RAND, AND RELATED SCHISTS
The Catalina Schist exhibits lithologic, metamorphic, and tectonic similarities with the Late Cretaceous Pelona, Orocopia, Rand, and related schists (PORS) of southern California and adjoining areas (Fig. 3.1; Haxel and Dillon, 1978; Ehlig, 1981; Jacobson and Sorenson, 1986, 2011; Grove et al., 2003b; Saleeby, 2003; Ducea et al., 2009; Chapman, 2016). The PORS are mostly younger than the Catalina Schist, but are similarly interpreted as trench (and potentially forearc-basin) materials emplaced beneath North American crust during low-angle subduction. It is thus reasonable to view the two as parts of a continuum. Both are distinguished from the Franciscan Complex by their inboard setting. The Catalina Schist and oldest part of the PORS were emplaced beneath the western belt of the Sierra Nevada–Peninsular Ranges arc, with younger PORS carried considerably farther beneath the continental edge. Both parts of the schist provide compelling evidence for low-angle subduction and tectonic erosion of the base of the overriding (North American) plate.
The older (northwestern) parts of the PORS underlie the western to medial portions of the southernmost Sierra Nevada batholith, analogous to the tectonic setting envisioned for emplacement of the Catalina Schist beneath the western edge of the northern Peninsular Ranges. Younger (central and southeastern) PORS bodies were carried farther beneath the margin, with the most inboard exposures located at least 300 km inland from the edge of the continent at the time of underplating (Jacobson et al., 2017).
The PORS are dominated by quartzofeldspathic metasandstone inferred to be derived from trench turbidite (Haxel and Dillon, 1978). The schists locally include up to 10% metabasite. Trace element compositions are mostly characteristic of normal and enriched mid-oceanic-ridge basalt, although a few samples exhibit alkaline within-plate or island-arc affinity (Haxel et al., 2002; Chapman, 2016). Metabasite is commonly associated with minor amounts of Fe- and Mn-rich quartzite interpreted as metamorphosed pelagic chert. Many of the schists also include rare lenses of serpentinite, up to hundreds of meters in length, encased within quartzofeldspathic schist. These bodies were presumably derived from dunite and peridotite. Relict dunite and peridotite (harzburgite), variably altered to serpentinite, have been recognized in one body of PORS (Cemetery Ridge; Haxel et al., 2015, 2018). Mineral and whole-rock compositions are most compatible with derivation from the overriding slab, but do not unequivocally rule out a source in the subducting plate. Centimeter- to meter-long pods of actinolite ± talc ± fuchsite within quartzofeldspathic schist are widespread throughout the PORS, although volumetrically minor. Relations at Cemetery Ridge indicate that actinolite formed by metasomatic alteration between ultramafic rock and host quartzofeldspathic schist (Haxel et al., 2015, 2018).
The PORS are pervasively recrystallized and characteristically exhibit inverted metamorphic field gradients (Haxel and Dillon, 1978; Jacobson et al., 1988; Chapman, 2016). Metamorphic assemblages range from epidote-blueschist and greenschist facies to upper-amphibolite or even lower granulite facies (Graham and Powell, 1984; Jacobson and Sorensen, 1986; Jacobson, 1995; Ducea et al., 2009; Chapman et al., 2011; Chapman, 2016). Mineral assemblages and textures most closely resemble those in the epidote-blueschist and epidote-amphibolite units of the Catalina Schist (Jacobson and Sorensen, 1986). Much of the quartzofeldspathic schist includes distinctive millimeter-scale porphyroblasts of albite or oligoclase colored gray to black by inclusions of graphite, a texture also observed in some parts of the Catalina Schist. In contrast to the Catalina Schist, true blueschists, as indicated by the presence of jadeite, aragonite, and/or lawsonite, have not been recognized in the PORS, but could be present at deep structural levels not yet exposed. Depth of metamorphism is not well constrained but is generally estimated to be in the range of 30–40 km (Chapman, 2016).
The relatively high-(P/T) metamorphism of the PORS and lithologic similarities with the Franciscan subduction complex have led most workers to interpret the schists as a relict subduction complex (Crowell, 1968; Yeats, 1968; Hamilton, 1988; Jacobson et al., 1996, 2011, 2017; Yin, 2002; Grove et al., 2003b; Saleeby, 2003; Chapman, 2016). Nonetheless, some workers have suggested that parts of the PORS were derived from the forearc basin (Hall, 1991; Barth and Schneiderman, 1996; Chapman, 2016), as has been proposed for the Catalina Schist (Grove et al., 2008).
Structurally, the PORS have the appearance of a coherent metamorphic terrane; i.e., they lack an obvious block-in-matrix character and individual lithologic bands can be traced for 100s of meters to kilometers along strike. Nonetheless, repetitive interlayering of metabasalt and metachert with metasandstone throughout the PORS is suggestive of tectonic mixing of oceanic crust and trench sediment (cf. Kimura et al., 1996). Multiple generations of highly attenuated isoclinal folds with axial-planar schistosity are widespread within the schist. The large ductile strains associated with this deformation would render initial fault boundaries difficult to recognize.
The age of the PORS protolith and time of underthrusting and exhumation are constrained by detrital zircon U-Pb ages and 40Ar/39Ar and related thermochronology (Grove et al., 2003b; Jacobson et al., 2007, 2011; Chapman, 2016). The oldest PORS, that in the San Emigdio Mountains of the southern tail of the Sierra Nevada, is at least 90 Ma in age, and may be as old as ca. 100 Ma (Chapman et al., 2013). The above dates indicate overlap with the protolith and metamorphic ages of the youngest, structurally deepest parts of the Catalina Schist (epidote-blueschist and lawsonite-blueschist and lower-grade units; Grove et al., 2008). The youngest PORS (Orocopia Schist and parts of the Pelona Schist) have protolith and metamorphic ages of ca. 70–65 Ma (Jacobson et al., 2017; Seymour et al., 2018). This part of the terrane postdates all magmatic activity in the Sierra Nevada–Mojave–Peninsular Ranges magmatic arc in California, and is coeval with active basement uplift and basin formation in the Rocky Mountains foreland region (DeCelles, 2004; Jones et al., 2011).
Although protolith and emplacement ages of the various bodies of PORS vary by 20–30 m.y. or more, the thermochronologic results indicate initial rapid cooling of each body immediately following underthrusting. This period of cooling is attributed to subduction refrigeration and/or erosional and tectonic denudation (Jacobson et al., 2007, 2017; Grove et al., 2019). A second cooling event, clearly associated with low-angle normal faulting, occurred during late Oligocene–early Miocene time; i.e., final exhumation of the PORS and Catalina Schist were driven by similar processes (Grove et al., 2008).
The detrital zircon ages indicate a systematic progression in sediment source area with time (Fig. 3.2). The schist in the San Emigdio Mountains is dominated by late Early Cretaceous ages characteristic of the western side of the Sierra Nevada–Peninsular Ranges arc (Grove et al., 2003b; Jacobson et al., 2011; Sharman et al., 2015). Minor Jurassic ages denote the earliest phases of arc magmatism. Paleozoic and Precambrian grains are probably derived from Lower Mesozoic arc wallrocks. The overall zircon age spectrum of the San Emigdio schist is remarkably like that of the similarly aged Catalina lawsonite-blueschist and lower-grade rocks. Younger parts of the PORS display progressively increasing contributions from the early Late Cretaceous eastern side of the Sierra Nevada–Peninsular Ranges arc and the more-inboard, latest Cretaceous Mojave arc and Proterozoic country rock.
We thank John Wakabayashi and Zeb Page for their helpful and constructive reviews, Josh Schwartz for editing, Gray Bebout for permission to use the map of the ultramafic mélange and for discussions, Sarah Penniston-Dorland and Mark Legg for discussions, the USC Wrigley Institute for logistical support, and the Catalina Island Conservancy for permission to access the field sites.
Figures & Tables
From the Islands to the Mountains: A 2020 View of Geologic Excursions in Southern California
This volume includes five geologic field-trip guides in the Los Angeles region associated with the 2020 GSA Cordilleran Section Meeting that was scheduled for May 2020, in Pasadena, California. The guides are organized in a generally counterclockwise order around the Los Angeles Basin. The first guide by Burgette et al. provides new slip rates, age constraints, and observations of the active Sierra Madre fault zone that borders the northern side of the San Gabriel and San Fernando Valleys. The Nourse et al. guide takes a new look at the San Gabriel Mountains from a basement and geomorphologic perspective. Further west, Keller et al. provide one of the first published field-trip guides focused on the 9 January 2018 Montecito debris flows that caused 23 deaths. The volume then moves south to Santa Cruz Island, where Davis et al. provide an updated review of the island’s geology within the California borderlands. The final guide returns to the east, where Platt et al. present the unique geology of Santa Catalina Island with a focus on the subduction-related Catalina Schist.