Recent advancements in geochronology, geologic mapping, and landslide characterization in basement rocks of the San Gabriel Mountains block
Published:May 18, 2020
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Jonathan A. Nourse, Brian J. Swanson, Alexander D. Lusk, Nicolas C. Barth, Joshua J. Schwartz, Karissa B. Vermillion, 2020. "Recent advancements in geochronology, geologic mapping, and landslide characterization in basement rocks of the San Gabriel Mountains block", From the Islands to the Mountains: A 2020 View of Geologic Excursions in Southern California, Richard V. Heermance, Joshua J. Schwartz
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This field trip examines Paleoproterozoic basement, Neoproterozoic metasedimentary strata, and crosscutting Mesozoic intrusive rocks at Frazier Mountain, Placerita Canyon, and Limerock Canyon in the western San Gabriel Mountains block, California. We present new U-Pb zircon geochronology results that constrain the Proterozoic through Cretaceous tectonic and magmatic history. The excursion ends in San Antonio Canyon in the eastern San Gabriel Mountains where several large rock avalanche deposits are sourced from distinct basement rocks. 10Be surface exposure ages and post-infrared infrared stimulated luminescence burial ages demonstrate late Pleistocene to Holocene movements for these landslides.
This field excursion visits spectacular exposures in the San Gabriel Mountains and related rocks of the San Gabriel Mountains basement complex near Frazier Mountain, where we will present new research results from three study areas (Fig. 0). We draw upon collective interests in field geology, petrology, and geochronology to describe relationships spanning Paleoproterozoic through Quaternary time. To reduce this article to a manageable size we present the large body of U-Pb zircon, 10Be surface exposure, and post-IR IRSL (post-infrared infrared stimulated luminescence) burial ages in a series of summary tables and plots. Rigorous discussion of the geochronology data, along with implications for tectonic reconstruction, etc. will appear in future publications (in preparation). However, there should be ample time in the field to discuss the data where we view key outcrops. This guidebook is divided into three primary parts based on research topics.
Part 1 of this guidebook, prepared by Brian Swanson, Alex Lusk, and Joshua Schwartz, describes Proterozoic and Cretaceous crystalline rocks south of the San Andreas fault (SAF) near Frazier Park at the west end of the San Gabriel Mountains block, as well as Cretaceous plutons of the Sierra Nevada batholith across the SAF to the north. New U-Pb dates of biotite gneiss, augen gneiss, deformed felsic dikes/sills, “splotchy” banded gneiss, and unfoliated granite on Frazier Mountain help inform the timing and history of gneiss formation and intrusive activity within the Proterozoic. We compare Cretaceous plutons of the Transverse Ranges on either side of Ridge Basin south of the SAF with Cretaceous plutons north of the SAF within the southern Sierra Nevada batholith and associated implications for evolution of the Pacific/North American Plate boundary. In addition, new mapping characterizes several large Neogene to Quaternary landslides that overprint these crystalline basement rocks on the eastern margin of Frazier Mountain.
Part 2, prepared by Jonathan Nourse and Karissa Vermillion, focuses on metasedimentary strata of the enigmatic Placerita Formation of Miller (1934) and Oakeshott (1958), exposed as pendants in Mesozoic plutons in the western San Gabriel Mountains. A Neoproterozoic depositional age is established for these continental margin sediments, and previously unrecognized Paleoproterozoic basement is identified. A series of distinctive Middle Jurassic and Late Jurassic intrusive rocks constrain a Late Jurassic age for the principal upper amphibolite facies metamorphic and deformational event. Late Cretaceous plutons in both areas record a culminating thermal event responsible for profound discordance of zircon data in the older rocks. A weak Laramide fabric is locally developed.
Part 3, prepared by Nicolas Barth, describes several large rock avalanche deposits in San Antonio Canyon of the eastern San Gabriel Mountains that are derived from distinct bedrock sources. Newly available lidar data and Quaternary geochronology lead to a paradigm-shifting view of the role of these landslides: (1) Large bedrock landslide deposits are three times as abundant as previously thought, covering ~25% of the San Antonio catchment by area. (2) Long-runout landslide deposits formerly mapped as early Quaternary (1–2 Ma) have been dated as late Holocene. These large landslides have driven considerable landscape changes including dammed lakes, major post-event aggradation, formation of epigenetic gorges, and reorganized drainages. Focus will also be placed on the identification of landslide deposits from geomorphology and internal rock textures.
PART 1. GEOLOGY AND U-Pb ZIRCON GEOCHRONOLOGY OF PROTEROZOIC AND LATE CRETACEOUS BASEMENT ON FRAZIER MOUNTAIN, WITH AN OVERVIEW OF LATE QUATERNARY LANDSLIDES (DAY 1, A.M. AND EARLY P.M.)
LOCATION AND SIGNIFICANCE
Frazier Mountain lies just south of the San Andreas fault (SAF) near the western end of the San Gabriel block in the central Transverse Ranges of southern California. The regionally significant San Gabriel fault intersects the SAF from the southeast, and the Garlock fault intersects the SAF from the northeast in a complex nexus on the northeast side of Frazier Mountain (Fig. 1). Proterozoic gneisses at Frazier Mountain and Cretaceous plutonic rocks exposed south of the SAF are juxtaposed against Cretaceous plutons of the Sierra Nevada Batholith on the north side of the SAF. Updated geologic mapping was conducted by the California Geological Survey in this area as part of the STATEMAP program in collaboration with the U.S. Geological Survey with the ultimate goal of creating a seamless geologic map of the 1:100,000-scale Lancaster 30′ × 60′ Quadrangle (Swanson and Olson, 2016; Olson and Swanson, 2017, 2019). New geochronologic and geochemical data obtained from the basement rocks during and following publication of these maps add to the understanding of Proterozoic metamorphism, magmatism, and deformation and refine the timing of Cretaceous plutonic activity in this area.
New mapping has also confirmed the presence of several large basement-involved landslides on the eastern flanks of Frazier Mountain and along the south side of the SAF to the east near Gorman. These slides record slope instability over multiple time periods beginning with ancient slides that appear to underlie and pre-date the Pliocene Hungry Valley Formation. An ancient mega-landslide complex southeast of Gorman lies on the Lower Member of the Hungry Valley Formation but originated from an unidentified source across the SAF to the north during deposition of the upper member of the formation. Other large landslides on the flanks of Frazier Mountain that overlap the Hungry Valley Formation were interpreted as Quaternary low-angle thrust faults or older alluvium by early mappers (e.g., Crowell, et al., 1982; 2002; Dibblee and Minch, 2006).
The geology of Frazier Mountain has been mapped at varying levels of detail, with the greatest detail focused on the eastern flank bordering Ridge Basin and on the western flank bordering Lockwood and Cuddy Valleys. Ridge Basin was mapped in detail by Crowell and his students and colleagues and published geologic mapping of basement rocks in adjoining areas has been conducted in modest detail. Crowell (1952a) published a geologic map of the Lebec Quadrangle, and Crowell (1954a) published a geologic map of the Ridge Basin area. Carman (1964) published early detailed descriptions for gneisses exposed on the west side of Frazier Mountain, and Burnett (1960) provides preliminary descriptions of exposures on the south side of the mountain. Crowell (1964) published early regional mapping of Frazier Mountain in general. Early mapping was compiled for the 1:250,000-scale Los Angeles Sheet of the Geologic Atlas of California by Jennings and Strand (1969). Dibblee (1974) subsequently produced an unpublished geologic map of the area and Weber (1988) focused his detailed mapping along the southern and eastern flanks of the mountain. Dibblee (1982a) provided an overview of the geology of Frazier Mountain and adjoining areas. Crowell et al. (1982, 2002) published geologic maps of Ridge Basin and the eastern flank of Frazier Mountain, and Crowell (2003a; 2003b) provided an overview of Ridge Basin and the rocks bordering the basin. Dibblee and Minch (2006) published 7.5′ quadrangle geologic maps of the Frazier Mountain and Lebec Quadrangles compiled from earlier mapping. Davis and Duebendorfer (1987) produced a detailed strip map of the geology along the SAF west of Tejon Pass and soils and geomorphic studies of the Frazier Mountain area were conducted by Crippen (1979) and Zhao (1989; 1990). Kellogg et al. (2008) published updated mapping of the western margin of Frazier Mountain, west of the Frazier Mountain Quadrangle. New mapping was conducted by the California Geological Survey as part of the STATEMAP project and the results of this mapping and associated geochronologic, geochemical and thin section analysis form the basis for new data and interpretations presented in this guidebook.
Frazier Mountain is underlain primarily by biotite gneiss, augen gneiss, “splotchy”-banded granulitic gneiss, and migmatitic gneiss, locally overprinted by mylonitic shear bands (Xgn1 on map figures). The biotite gneiss varies from dark gray and white with cm-scale banding to dominantly white with splotchy banding with local granulitic texture and varying biotite content on the east side of Frazier Mountain. Quartzite, metaconglomerate, and marble are rare constituents. The biotite gneiss has previously been interpreted as a paragneiss resulting from metamorphism of dominantly pelitic sediments and local interleaved volcanic beds (Silver, 1971; Frizzell and Powell, 1982), although sillimanite was not found to be a prominent constituent. Silver et al. (1963) and Silver (1971) correlated this unit with the Paleoproterozoic (1750–1680 Ma) sedimentary assemblage identified elsewhere in the Transverse Ranges at Soledad Basin and the western San Gabriel Mountains. Augen gneiss, likely correlative to the Soledad gneiss containing discrete, lens-shaped to nearly euhedral potassium feldspar grains up to 6 cm long forms large mappable bodies on the western and northern sides of the mountain (Xagn and Xagn-cat on map figures). Alluvial debris indicates that additional unmapped bodies are likely present on the northeastern side of the mountain. Carman (1964) interpreted the augen to be metasomatic in origin, but more recent studies conclude they were originally magmatic megacrysts within a porphyritic granodiorite or quartz diorite that intruded the surrounding biotite gneiss in Proterozoic time and have subsequently been overprinted by metamorphism and deformation, including development of foliation and spatially variable mylonitic shearing (e.g., Silver, 1971; Frizzell and Powell, 1982). Premo in Kellogg (2001) determined a Paleoproterozoic U-Pb age of 1690 ± 5 Ma for the augen gneiss exposed southwest of Frazier Mountain. New U-Pb dates obtained for the Proterozoic Rocks at Frazier Mountain are presented herein.
Previous studies have shown a strong correlation between gneissic rocks exposed at Frazier Mountain, southwest of the San Gabriel fault, with similar rocks exposed northeast of the fault in the Soledad Basin and San Gabriel Mountains (e.g., Silver, 1966, Ehlig, 1981; Frizzell and Powell, 1982). Such correlations support ~40–60 km of long-term, right-lateral displacement on the San Gabriel fault between ca. 12 and 5 Ma (e.g., Crowell, 1952b, 1954b, 1982, 2003c; Ehlig, 1975b; Ehlig and Crowell, 1982; Premo et al., 2007). Correlations reported for younger overlying units such as similar Oligocene volcanic rocks of the Plush Ranch and Vasquez Formations (e.g., Carman, 1964; Frizzell and Weigand, 1993), and for the Miocene Caliente and Mint Canyon Formations (Ehlert, 1982, 2003; Hoyt et al., 2018) indicate similar magnitude right-lateral slip on the San Gabriel fault. All of these units have been interpreted to correlate with similar rocks and source areas offset again across the SAF in the eastern Transverse Ranges to the southeast (e.g., Crowell, 1954b, 1962; Ehlert, 1982, 2003; Powell, 1993; Dickinson, 1996; Nourse, 2002). In particular, Powell (1993) correlated the gneissic rock assemblage at Frazier Mountain with the Hexie Mountain assemblage in the eastern Transverse Ranges, which includes the distinctive augen gneiss of Monument Mountain.
Cretaceous plutonic rocks and associated pendant rocks form the exposed crystalline basement rock north of the SAF and bordering the east margin of Ridge Basin. Smaller unnamed Cretaceous bodies are exposed near the northwestern and southeastern flanks of Frazier Mountain (Kg1 and Kg2 on map figures, respectively). Crowell (1952a) first defined and named the Tejon Lookout Granite (Ktl) for exposures north of Gorman and the SAF, the Lebec Granodiorite (Kle) for exposure north of the Garlock fault near the town of Lebec, and the Liebre Granodiorite (quartz diorite) (Kli) exposed at Liebre and Bald Mountains to the east of Ridge Basin and south of the SAF. Ross (1972, Ross 1989) published petrographic descriptions and geochemical data for these rocks. Updated petrographic descriptions and U-Pb dates for these rocks are provided in this guide. The pendant rocks enveloped within the plutons north of the SAF are presumed to be Paleozoic or possibly early Mesozoic in age but have not been directly dated. The Cretaceous plutonic rocks are thought to represent intrusion of magma associated with subduction of the Farallon Plate below the western margin of North America. These plutons have been offset by the SAF and rotated within the Transverse Ranges during late Tertiary evolution of the plate boundary (e.g., Nicholson et al., 1994; Crouch and Suppe, 1993). In general, Cretaceous plutonic rocks in the Transverse Ranges appear to be slightly younger than along strike rocks to the north in the Sierra Nevada batholith (e.g., Nadin et al., 2016) and to the south in the northern Peninsular Ranges (e.g., Premo et al. 2014).
Buwalda et al. (1930) first interpreted the presence of a very low-angle thrust fault on the south flank of Frazier Mountain, termed the Frazier Mountain fault. The fault was interpreted based on a lobe of gneiss found overlapping the Pliocene Hungry Valley Formation along a gently dipping contact and formed the basis for future interpretations of the geologic structure underlying Frazier Mountain (e.g., Crowell, 1950, 1982; Kellogg, 2004; Dibblee and Minch, 2006). Detailed mapping by Weber (1988), however, recognized that the lobe of gneiss exhibits a brecciated texture typical of rock avalanche debris and the lobe actually represents an old landslide, and subsequent maps by Crowell et al. (2002) and Swanson and Olson (2016) follow this interpretation. Similarly, several slivers of basement rock northeast of Frazier Mountain along the south side of the SAF were originally interpreted as thrust fault slivers flowering off of the SAF (e.g., Crowell et al., 1982). Subsequent mapping has also reinterpreted these features as old landslides (Weber, 1988; Crowell, et al., 2002; Swanson and Olson, 2016; Olson and Swanson, 2017). Weber also mapped a large ancient landslide complex that overprinted the basement rock on the northeast flank of Frazier Mountain and interpreted the age of movement of several slides in this area as pre-dating deposition of the Hungry Valley Formation. These landslides will be a topic of further discussion on this field trip.
GEOLOGIC UNITS, PETROGRAPHIC DESCRIPTIONS, AND MAP RELATIONS
Proterozoic Basement Rocks of Frazier Mountain
The distribution of geologic units discussed herein and the locations of most U-Pb samples are shown on the geologic maps in Figures 2 and 3, which were extracted and modified from Swanson and Olson (2016). Paragneiss mapped as Xgn1 on the map includes a variety of textures, with foliated biotite gneiss being the primary facies exposed at Frazier Mountain. In outcrop, the paragneiss foliation is defined by gray, light-gray, and pinkish-gray compositional banding on the mm to cm scale. The gneiss is locally migmatitic with felsic melt bands. Banding is locally variable in intensity and lateral continuity. Gneiss with a granulitic texture and distinctive splotchy banding is a common, but currently undifferentiated, subunit exposed east of the summit of Frazier Mountain. The second most abundant Proterozoic rock at Frazier Mountain is augen orthogneiss (Xagn), which contains ovoid to nearly euhedral potassium feldspar porphyroclasts and is interpreted as having a porphyritic granodiorite protolith. Mylonitic textures locally overprint the foliated paragneiss and augen gneiss, particularly near the summit of Frazier Mountain. Some bands of augen gneiss are also locally overprinted by cataclastic deformation and are differentiated as Xagn-cat on the map figures. Alaskite sills/dikes are locally pervasive and typically follow the foliation in both the banded paragneiss and the augen gneiss. Small pods of undeformed Proterozoic granite were found to intrude the augen gneiss on the southwest side of Frazier Mountain (Xgr). Southeast of Frazier Mountain and the Frazier Mountain fault, isolated outcrops of paragneiss associated with Cretaceous granite are more schistose in texture and contain more abundant garnet (Xgn2). Bulk-rock major oxide geochemistry confirms a quartzofeldspathic composition for all of the Proterozoic units.
Foliated biotite paragneiss underlies much of Frazier Mountain (Fig. 4A). At the microscale, paragneisses preserve a variety of textures ranging from foliated to coarse-grained granoblastic (see splotchy-banded gneiss below) to locally mylonitic. A general composition comprises quartz (35–50%), potassium feldspar primarily as microcline with tartan plaid twinning (15–20%), plagioclase feldspar, which is commonly retrogressed (20–35%), biotite (5–10%), and <5% each of titanite and opaques. Accessory minerals include zircon and apatite. Secondary white mica (sericite) and chlorite form from reaction of plagioclase and biotite, respectively. In rocks that preserve a gneissic texture, quartz is medium to coarse grained and subhedral to ameboid in shape. Biotite and white micas have irregular shapes, ranging from blocky books to elongate; alignment of these phases helps to define a moderately strong to strong macroscopic foliation. Gneissic banding is defined primarily by higher biotite content of dark bands and higher quartz and feldspar content of light bands. A mylonitic foliation, where present, is commonly defined by both the alignment of phyllosilicates and a shape-preferred orientation of recrystallized or deformed quartz grains (discussed in more detail below). A variant of biotite gneiss with more schistose foliation, a slightly younger U-Pb date, and cut by Cretaceous granite crops out locally southeast of the Frazier Mountain fault and is referred to later as Xgn2 (Fig. 4B).
Augen gneisses are distinctive from other gneiss facies by ubiquitous mm- to cm-scale ovoid to nearly euhedral potassium feldspar porphyroclasts (augen) wrapped by a foliated medium- to dark-gray biotite-rich gneissic matrix (Fig. 5A). Augen are pink to white perthitic microcline interpreted to be inherited from a porphyritic granodioritic or quartz monzonitic protolith that intruded the Paleoproterozoic quartzofeldspathic paragneisses (Xgn). Long axes of the augen appear variably oriented in weakly foliated exposures, but tend to be rotated or stretched parallel to the macroscopic foliation in more intensely foliated exposures. Streaky lineation is locally apparent, particularly where overprinted by mylonitic shearing. The augen gneiss is also locally overprinted by brittle cataclastic deformation in local bands (Xagn-cat).
Petrographic observation reveals a composition comprising potassium feldspar augen in a variably proportioned matrix of plagioclase, quartz, potassium feldspar, and biotite, with minor zircon, opaques, and apatite. Microcline augen are found to be either single crystals or aggregates. Potassium feldspar in the matrix is generally also microcline with tartan plaid twinning and is sub- to anhedral, forming grains subordinate to other phases. Plagioclase, where not retrogressed, is anhedral, has limited polysynthetic twinning, and is locally associated with minor myrmekite. Quartz forms globular grains with pervasive undulatory extinction and subgrain development. Biotite forms coarse, blocky grains which define the primary gneissic foliation and compositional banding. Biotite is also found as primary inclusions within plagioclase.
Augen gneisses record moderate to extensive retrogression of plagioclase and biotite. Plagioclase records variable intensities of retrograde reactions to sericite or epidote group minerals (e.g., epidote, zoisite, clinozoisite). The original composition of the feldspar may be the primary control on reaction products; zoisite and clinozoisite may react from calcic plagioclase whereas sericite comes from sodic endmembers. Where fractured, chlorite commonly forms at the expense of biotite. The abundance of fluid-bearing reaction products indicates that fluids likely played an important role in retrogression.
Alaskite Sills and Dikes (within Xgn1 Paragneiss and Augen Gneiss)
A series of leucocratic granitoids locally occur both within the augen orthogneiss and paragneiss as elongate sills and dikes 5–20 cm thick (see Figs. 5B and 5C). Most sills are parallel to the gneissic foliation, but some crosscut the foliation. In outcrop, the alaskite is fine to medium grained, with potassium feldspar polycrystalline aggregates sometimes exceeding 1 cm in diameter. These rocks weather to a pale yellowish-brown to light-gray color. Contacts with the host gneiss tend to be well defined and distinct although minor host material is sometimes incorporated into the alaskite matrix. The genetic relationship between alaskite and the surrounding host rocks remains unclear.
The alaskite preserves a hypidiomorphic-granular texture comprising plagioclase feldspar (30–40%), quartz (25–35%), potassium feldspar primarily observed as microcline aggregates with tartan plaid twinning and exsolution lamellae (25–35%), biotite (5–10%), and epidote group minerals (<5%), along with trace apatite, zircon, titanite, and opaques. Mafic clots are commonly associated with incorporated orthogneissic host material and are rare within the igneous matrix. Epidote group minerals, most commonly associated with mafic clots, have rare metamict cores (possibly related to the presence of allanite).
In thin section, alaskites are largely undeformed internally with only limited solid-state deformation evidenced by pervasive subgrain development and rare dynamic recrystallization along quartz grain boundaries. Feldspars preserve rare myrmekite and are commonly reacted to sericite and epidote group minerals, indicating probable fluid infiltration.
Splotchy-Banded Granulite Gneiss (within Xgn1 on Fig. 3)
A distinctive facies of Xgn1 is characterized by a leucocratic matrix with dark, cm-scale thick, splotchy, and mottled discontinuous mafic bands defining a distinct but discontinuous foliation (see Fig. 7A). Exposure of the splotchy banded quartzofeldspathic gneiss is generally limited to areas east of the summit of Frazier Mountain, south of the primary trace of the SAF, but has not yet been differentiated by mapping.
Micro-scale textures reflect what is observed in outcrop. The microstructure is dominated by a medium- to coarse-grained quartzofeldspathic matrix cut by splotchy, dark bands. The matrix is dominated by granoblastic quartz, plagioclase, and microcline, with rare white mica and accessory minerals including abundant zircon. Splotchy bands are made up of layered mafic clots comprising fine- to medium-grained biotite, medium-grained plagioclase and potassium feldspar, fine-grained granoblastic quartz, amphibole (hornblende), titanite, opaques, and rare epidote group minerals. Ilmenite is inferred based on polycrystalline titanite rims. Biotite and hornblende (where present) have bimodal grain size distributions: both form fine-grained, equant, stubby aggregates and, less commonly, elongate laths. The macroscopic splotchy foliation is defined compositionally, with only minor alignment of elongate grains at the micro-scale. Feldspar is present to the extent that it tends to form interconnected networks which may form a stronger framework than the surrounding quartz and micas. This gives rise to crystal-plastic deformation within the feldspar, which is discussed in further detail below.
More mafic facies lack the macro-scale discontinuous foliation defined by compositional banding. At the micro-scale, they are compositionally and texturally similar to the mafic bands within the quartzofeldspathic rocks. The mafic facies are made up chiefly of plagioclase, biotite, hornblende, minor quartz, titanite, and epidote group minerals. No potassium feldspar is observed in the mafic facies. Secondary alteration is limited to minor sericitization of plagioclase feldspar and minor reaction of biotite to chlorite. Minor carbonate and quartz are present in veins.
Plagioclase and microcline grains generally form an interconnected load-bearing framework since they tend to be stronger than the surrounding phases. This is especially prevalent in the quartzofeldspathic layers. Within these load-bearing frameworks, feldspar records variable intensities of dynamic recrystallization, chiefly by bulge nucleation, as evidenced by sutured or irregular grain boundaries and adjacent small recrystallized grains. Recrystallized zones are generally a few grains thick and concentrated along grain boundaries between relict grains, but also form recrystallized “necklaces” that cut through relict grains. Recrystallization of feldspar indicates approximate temperature >450 °C.
Quartz records evidence for crystal plastic deformation in the form of extensive development of subgrains, undulose extinction, and moderate dynamic recrystallization by subgrain rotation. Approximately equant recrystallized grains mantle larger gneissic grains in which subgrains and undulose extinction are developed. Recrystallization in quartz is not homogenous throughout the microstructure; this is likely because the feldspar framework armored some quartz grains while others were preferentially deformed.
Granite (Map Unit Xgr on Fig. 2)
Undeformed granitoids crop out as isolated, 100-m-scale bodies within the augen gneiss (Xagn) on the southwest side of Frazier Mountain (see Fig. 2). These bodies were found to cut the primary gneissic foliation and locally envelope 10-cm- to 1-m-scale blocks of Xagn (see Fig. 7B). In outcrop, the undeformed granitoid is medium to fine grained and has a light-gray, speckled appearance. Aligned mafic grains and rare cm-scale-thick compositional bands define a very weak and variable magmatic foliation; no solid-state deformation has been observed within these bodies. Weathered rock maintains its competence with weathered surfaces forming a thin, reddish-brown rind. Bulk-rock major-oxide chemistry classifies this unit as an alkali-rich granite or quartz monzonite, per the Middlemost (1994) total alkali versus silica (TAS) discrimination scheme (Fig. 6).
Compositionally, the undeformed granitoid records a magmatic texture with plagioclase (50–55%), quartz (15–20%), microcline (15–20%), biotite (10–15%), and less than 5% each of epidote (including allanite), titanite, and white mica (mainly retrogressive from at the expense of plagioclase but also occasionally interleaved in biotite), as well as trace zircon, apatite, and opaque minerals. Both feldspars form relatively large, up to 2 mm poikilitic grains; inclusions commonly include feldspar, quartz, and apatite, all of which are anhedral and rounded. Biotite, epidote group minerals, anhedral titanite, and opaques tend to form intergrown clusters. Allanite commonly forms the cores of epidote group grains.
Plagioclase has been moderately to extensively reacted to sericite and also records rare myrmekite along grain boundaries. Although not pervasive, joints concentrate retrogressive minerals suggesting a history of fluid infiltration and reaction, consistent with reactions recorded in plagioclase breakdown and metasomatism. However, nowhere is biotite observed to breakdown to chlorite.
At the outcrop and hand-specimen scales, there is no evidence for solid-state deformation. Nonetheless, at the microscopic scale, quartz grains contain subgrain boundaries and undulose extinction, indicative of limited crystal plastic deformation. Quartz grain shapes remain unstrained.
Structures Associated with the Gneissic Rocks
Where present, gneissic banding is often folded into cm- to m-scale open to isoclinal folds. Paragneisses commonly record evidence of crystal-plastic and local brittle deformation. In outcrop, fractures are common but only record limited displacement. Quartz, calcite, white mica, epidote group minerals, and opaques fill fractures and suggest infiltration of fluids may have been instrumental in secondary mineralization. Fluid infiltration is corroborated by extensive and often complete sericitization of plagioclase feldspar.
Protomylonitic to mylonitic textures record variable intensities of dynamic recrystallization and grain-size reduction in high-strain rocks. In areas, quartz forms ribbons indicative of easy glide and preserves moderate to extensive dynamic recrystallization, primarily by subgrain rotation and bulge nucleation. Gneissic quartz grains show varying intensities of internal lattice strain and buildup of dislocations as evidenced by undulatory extinction and development of subgrains, similar in size and shape to recrystallized grains. Within protomylonites and mylonites, feldspars show only limited evidence for strain accommodation by crystal-plastic processes. In most samples, sericite has formed at the expense of plagioclase feldspar and, in mylonitic rocks, appears to accommodate much of the strain, often forming elongate bands parallel to the mylonitic foliation. Together with quartz ribbons, these two structures define the primary mylonitic foliation.
Augen gneisses record evidence of both brittle and ductile deformation. Cataclastic augen gneiss shows extensive fracture at the grain scale, resulting in a brecciated texture in outcrop. Fractures commonly show reactions to sericite and chlorite in adjacent grains and are filled with precipitated quartz. Ductile deformation is evidenced by the extensive buildup of crystal-plastic strain in quartz and limited dynamic recrystallization. Recrystallization often manifests as necklaces of new recrystallized grains that cut through relict grains as well as producing sutured (bulging) and serrated grain boundaries. Overall, dynamic recrystallization is limited and where present, primarily occurs by bulge nucleation and subgrain rotation.
Pervasive mylonitization is spatially limited to outcrops of gneiss near the summit of Frazier Mountain and southward toward the Frazier Mountain fault; mylonites decrease in abundance west of the summit and were not observed in the splotchy-banded gneiss east of the summit, or in the middle Proterozoic granite (Xgr). In general, mylonitic rocks show extensive reduction in grain size compared to their respective protoliths (Xgn1, Xagn) and have a moderate to strong mylonitic foliation. They share similar compositions but consistently show more extensive retrogression of plagioclase feldspar to sericite and biotite to chlorite. These phyllosilicate reaction products form elongate, high-strain bands, evidenced by local shear bands within the phyllosilicates. Together with ribbon quartz, these phyllosilicate bands define the macroscopic mylonitic foliation. Quartz and relict feldspar can often be seen floating in an otherwise interconnected network of fine-grained phyllosilicates. Within these high-strain phyllosilicate zones are darker bands that likely form as a result of dissolution of soluble phases, leaving behind µm-scale thick dark bands of concentrated insoluble phases. The presence of these dissolution bands indicates that pressure solution, or another form of solution-reprecipitation creep, was an active deformation mechanism.
Observations at the micro-scale clearly indicate that dislocation creep was also active. Quartz records evidence for extensive crystal-plastic behavior; gneissic grains form elongate ribbons which have (1) sweeping undulose extinction, (2) development of subgrains that are generally also elongate in shape and subparallel to the ribbons, (3) deformation lamellae, and (4) variable intensities of dynamic recrystallization by both bulge nucleation and subgrain rotation.
Evidence for brittle behavior is also prevalent throughout the mylonitic rocks. Quartz, although primarily deformed by dislocation creep, is sometimes fractured, likely recording continued deformation to lower temperatures. Where not reacted to sericite, plagioclase feldspar often shows brittle fracture and kinking. In outcrop, microcline augen are elongate and have tails suggestive of ductile behavior; upon examination in thin section, augen are pervasively fractured with moderate dynamic recrystallization along these fractures. Where recrystallized, the microcline appears to chemically change to more sodic or calcic compositions.
Based on the active deformation mechanisms in quartz (dislocation creep with recrystallization by subgrain rotation, bulge nucleation, and brittle fracture) deformation temperatures likely range from ~400 °C to below the brittle-ductile transition (~300 °C) (Stipp et al., 2002).
Cretaceous Intrusive Rocks of Transverse Ranges, South of the SAF
Granite Northwest of Frazier Mountain (Kg1)
Light pinkish-gray, two-mica granite (Kg1) is exposed south of the SAF on the northwest flank of Frazier Mountain, on the northwest side of the Cuddy Valley fault. The granite is highly brecciated, suggesting possible overprinting by landslide movement, and in thin section, the rock is pervasively pulverized with all phases extensively fractured. Kg1 was previously mapped as a leucocratic gneiss by Carman (1964).
Plagioclase cores commonly record moderate reaction to sericite and quartz grains have variably sutured grain boundaries, moderate subgrain development, and sweeping undulose extinction. Minor myrmekite is also present, mostly found on the margins of larger plagioclase grains.
Granite Southeast of Frazier Mountain (Kg2)
Granitic bodies crop out southeast of Frazier Mountain on the southeast side of the Frazier Mountain fault. Kg2 intrudes a biotite gneiss/schist (Xgn2) and is unconformably overlain by the dominantly Pliocene Hungry Valley Formation (Thvl). In outcrop, granites are leucocratic, speckled with small (<1 mm), aligned biotite flakes that define a moderate to strong magmatic foliation. Weathered surfaces take on a yellowish-gray color. Geochemical analyses are currently in progress
This unit has a fine- to medium-grained, hypidiomorphic texture comprising plagioclase (35–45%), quartz (25–35%), potassium feldspar (25–35%), biotite (<5%), epidote group minerals (likely zoisite; <5%), trace primary and secondary white mica and chlorite, and zircon. Potassium feldspar commonly has small, rounded quartz and plagioclase inclusions and forms variably sized, irregularly shaped grains subordinate to the plagioclase. Grains commonly show tartan plaid twinning. Plagioclase, in some places with polysynthetic twinning, forms sub- to anhedral grains with diameters up to ~2 mm. Secondary alteration is evidenced by rare myrmekite, reaction of plagioclase to white mica, and limited reaction of biotite to chlorite. Quartz grains show limited subgrain development and undulose extinction.
This intrusive was first described by Crowell (1952a) as the Liebre Quartz Monzonite, although subsequent work indicates that granodiorite is more representative of the unit overall. Exposure of the Liebre Granodiorite is limited to the south side of the SAF although several isolated slivers mapped along the southern margin of the fault zone are also tentatively correlated with the Liebre Granodiorite based on similarities in composition and U-Pb zircon ages. Current mapping and previous work by Sexton (1990) reveal that the Liebre Granodiorite typically consists of speckled, light- to medium-gray, medium- to coarse-grained, generally equiangular, hypidiomorphic texture. In outcrop, Kli bodies are cut by aplitic to pegmatitic dikes up to several feet in thickness and several hundred feet in lateral extent and a weak magmatic foliation defined primarily by aligned mafic minerals is variably expressed (Sexton, 1990). More-mafic enclaves, generally cm-scale, ellipsoidal, and dioritic in composition, as well as schlieren, are present throughout but with significant spatial variability in abundance. In natural outcrops, the Kli commonly is moderately to highly decomposed to grus, with more mafic facies particularly susceptible to extensive weathering. Felsic dikes are more resistant and are more likely to form outcrops.
At the microscopic scale, Liebre Granodiorite records a magmatic texture ranging from allotriomorphic- to hypidiomorphic-granular. Primary constituents consist of plagioclase feldspar (55–65%), quartz (20–25%), biotite (10–20%), potassium feldspar (10–20%), minor (< 5% each) hornblende, titanite, epidote group, and opaques, plus accessory zircon and apatite. Plagioclase grains are hypidiomorphic to idiomorphic, generally sodic in composition commonly with polysynthetic twinning and oscillatory zoning. Plagioclase interiors tend to have higher concentrations of inclusions, chiefly biotite. Biotite in the matrix forms books up to 5–8 mm in maximum dimension. Potassium feldspar is anhedral, forming large (up to 1 cm) poikilitic grains that have interstitial textures. Quartz also tends to be anhedral with grains subordinate to plagioclase and biotite. Biotite and hornblende form mafic clots which are also commonly associated with epidote group minerals. Although much of the epidote is secondary and fills fractures, isolated subhedral grains up to >500 μm may be of magmatic origin. Titanite grains, which range from euhedral diamond-shaped to anhedral, are reportedly up to 2 mm in length (Sexton, 1990). Titanite is often associated with opaque minerals which have previously been reported as magnetite (Ross, 1972) but may also include ilmenite. Both zircon and apatite are present throughout as accessory phases.
Liebre Granodiorite records minor secondary alteration primarily in the form of veining and retrogressive mineral reactions. Myrmekite and fractures filled with epidote group minerals and quartz indicate reaction and precipitation of material by fluid phases. Biotite and calcic plagioclase both show minor to moderate reaction to chlorite and sericite, respectively; these reaction products are also consistent with fluid interaction. Hornblende, where present, tends to be partially reacted to biotite.
Pervasive fracturing of quartz and feldspar is common in Kli slivers that crop out within the SAF zone. Within these fault-zone rocks, fractures are mostly free of any secondary fill or precipitated material.
Quartz records evidence for crystal-plastic deformation in the form of moderate to extensive subgrain development and sweeping undulose extinction, occasionally manifesting as chessboard extinction. Minor dynamic recrystallization occurs within quartz, typically concentrated in thin zones along grain boundaries. Biotite grains show rare kinking. Ductile deformation is likely of minor importance in terms of finite strain and has not resulted in the development of a macroscopic solid-state texture.
Cretaceous Intrusive Rocks of Sierra Nevada Batholith, North of SAF
The Tejon Lookout Granite, as first named by Crowell (1952a) and later referred to as the Granite of Tejon Lookout by Ross (1972, 1989), comprises several facies with varying compositions and textures. Volumetrically most abundant is a medium-grained, light-gray to yellow-brown granite; the remainder comprises fine- to medium-grained biotite + hornblende granodiorite, aplitic and pegmatitic bodies, and minor mafic segregations. Good outcrops are rare as it tends to be deeply weathered along pervasive joints and sheared fractures. Exposure of Ktl is limited to large bodies north of the SAF and as small slivers or lenses within the fault zone. Bulk-rock major oxide geochemical analyses on confirm granitic and granodioritic compositions, per the Middlemost (1994) total alkali versus silica (TAS) discrimination scheme (Fig. 6).
A generalized granitic facies composition comprises variable proportions of potassium (35–40%) and plagioclase (20–25%) feldspars, quartz (30–35%), biotite (5–10%), white mica (<5%), and accessory titanite, zircon, and opaques (<5%). Potassium feldspar (chiefly microcline) forms both intergranular anhedral grains and large (up to 1 cm) phenocrysts commonly with perthitic twinning and abundant inclusions including quartz, plagioclase, biotite, white mica, and zircon. Where present, quartz inclusions in feldspar tend to be equant with rounded grain boundaries and define rare growth zoning as concentric, inclusion-rich rings. Grain size varies within the unit from medium- to coarse-grained. Potassium feldspar is generally coarser-grained than plagioclase with plagioclase commonly forming overgrowth rims around microcline; both show Carlsbad twinning.
The granodioritic facies comprises an assemblage of plagioclase (45–55%), amphibole (hornblende) (15–25%), biotite (15–20%), potassium feldspar (10–20%), (quartz (10–20%), and accessory epidote group, titanite, zircon, apatite, and opaques (<5% total). This facies tends to be fine grained with phenocrysts of hornblende and plagioclase. Plagioclase is commonly zoned, forming subhedral to euhedral tabular to equant grains, with rare larger phenocrysts. Quartz and potassium feldspar both fill interstitial spaces and have anhedral, globular grain shapes, sometimes completely enveloping other plagioclase, hornblende, or biotite grains. Tabular biotite grains commonly include intergrown chlorite and epidote group minerals (?), likely recording retrogression of biotite. Although largely preserving a magmatic texture, quartz records limited, and variable internal lattice strain evidenced by development of subgrains and sweeping undulatory extinction.
Where exposed in zones of fracturing or jointing, Ktl records breakdown reactions indicative of retrogression and/or alteration by fluids. Where retrogressed, chlorite and epidote group minerals grow mimetically at the expense of biotite; plagioclase, especially in zoned cores, is moderately to extensively sericitized. Minor myrmekite throughout the unit is consistent with fluid interaction. Proximal to mapped faults, moderate to intense fracturing is common. At the grain scale, these manifest as tensional fractures, variably filled with quartz or epidote; at the outcrop scale, fractured rocks are grusified and friable.
Cretaceous Lebec Granodiorite North of Garlock Fault (Map Unit Kle on Fig. 3)
In outcrop, the Lebec Granodiorite (also referred to as the Lebec Quartz Monzonite by Crowell, 1952a) presents as light-gray to peppery, medium- to coarse-grained biotite granodiorite. It is locally porphyritic with potassium feldspar phenocrysts and can be distinguished by rounded gray quartz masses and brown-stained mafic minerals. Crowell (1952a) reports two conspicuous facies: a light-gray, medium-grained unit and a pale-yellowish-brown coarser-grained unit. The yellowish-brown coloration is possibly related to secondary limonite staining. Exposure is limited to south of the Pastoria fault, north of the SAF, and northwest of the Garlock fault.
Bulk-rock major oxide and trace-element chemistry were measured for samples P-72, WP-433a, and WP-517. Samples WP-433a and P-72 both classify as granodiorite per the Middlemost (1994) total Alkali versus silica (TAS) discrimination scheme (Fig. 6). A slight increase in both alkalis and silica result in a granitic classification for sample WP-517 (possible Ktl), using the same classification scheme.
In thin section, primary magmatic textures are largely preserved. The Lebec Granodiorite is generally composed of variably proportioned potassium and plagioclase feldspar (50–60%), quartz (20–30%), biotite (10–20%), white mica (<5%), titanite (<5%), epidote group minerals (<5%), and accessory zircon and apatite. Rare and very minor garnets are present in some samples. Plagioclase feldspars are commonly albitic, evidenced by polysynthetic twinning, often show oscillatory zoning, and are subhedral. Potassium-rich feldspars are anhedral to subhedral, filling interstitial space, and commonly poikilitic, even preserving inclusion-rich concentric zonation rings. Both perthite and tartan plaid twinning are common. Quartz is anhedral to globular, forms interstitial to other phases, and is commonly found as inclusions in potassium feldspar grains. Biotite and white micas have irregular shapes, ranging from blocky books to elongate. In limited samples, biotite has a reddish-brown color suggesting high Ti content (and high temperature), consistent with variably proportioned but pervasive magmatic titanite. Hornblende with colorless clinopyroxene cores is reported as a minor but widespread constituent (Ross, 1972); we were unable to confirm the presence of either of these phases based on our sampling and petrography.
Where retrogressed, plagioclase is moderately to pervasively sericitized. Where zoned, the plagioclase cores commonly record more complete sericitization compared to the rims. Chlorite and minor epidote grow pseudomorphically at the expense of biotite. Secondary calcite is present in some samples, limited to within and around fractures.
Limited solid-state deformation is recorded by development of subgrains and sweeping undulose extinction in quartz. Evidence for solid-state deformation is generally associated with more extensively retrogressed samples. Occasional myrmekite suggests later interaction with metasomatic fluids.
U-Pb ZIRCON GEOCHRONOLOGY
Sampling Strategy and Analytical Methods
Eighteen samples were collected from crystalline basement rocks in the Frazier Mountain and Lebec Quadrangles between 2015 and 2018 for U-Pb zircon dating and whole-rock geochemical analyses. Eight of the samples were collected from known or suspected Proterozoic rocks at Frazier Mountain (see map Figs. 2 and 3), and two additional samples were collected from undated but presumed Cretaceous plutons near the northwestern and southeastern flanks of Frazier Mountain. Four samples were collected from the Lebec Granodiorite and Tejon Lookout Granite to the north of the SAF, one from the Liebre Granodiorite south of the SAF, and three additional samples were collected from granitoids of uncertain affinity along the SAF (see map Figs. 3, 8, and 9). Chapman and Saleeby (2012) and Chapman et al. (2012) published an age range of 88–92 Ma for the Lebec Granodiorite to the west of the study area, and the sample site for the current study area was selected to assess the age of granitoid rocks exposed between the main and north branches of the Garlock fault. The three samples from Tejon Lookout Granite targeted three facies observed within the pluton; no previous published ages were found for this granite. Kistler et al. (1973) reported an age of ca. 110 Ma for the Liebre Granodiorite from exposures east of the study area based on Sr87/Sr86 ratios.
Zircons were separated following standard methods involving crushing, pulverizing with jaw crusher and disk mill, and density separation on a Wilfley gold table and with heavy liquids. A subset of the jaw-crushed material was reserved for whole-rock geochemistry. These samples were powdered in an alumina ceramic shatter box and major- and trace-element analyses were conducted at Pomona College. Disk milled material was processed through a Frantz isodynamic separator (side tilt = 5°, front tilt = 20°) at 1.5 A to remove magnetic (non-zircon) minerals. Zircons were poured onto double-sided tape and mounted in epoxy, ground, and polished. They were then imaged on a Gatan MiniCL detector attached to an FEI Quanta 600 scanning electron microscope at California State University Northridge.
Uranium-lead ratios were collected using a ThermoScientific Element2 sector field–inductively coupled plasma mass spectrometry coupled with a Teledyne Cetec Analyte G2 excimer laser (operating at a wavelength of 193 nm). Prior to analysis the Element2 was tuned using the NIST 612 glass standard to optimize signal intensity and stability. Laser beam diameter was ~25 microns at 10 Hz and 75%–100% power. Ablation was performed in a HelEx II Active 2-Volume Cell™ and sample aerosol was transported with He carrier gas through Teflon-lined tubing, where it was mixed with Ar gas before introduction to the plasma torch. Flow rates for Ar and He gases were as follows: Ar cooling gas (16.0 NL/min), Ar auxiliary gas (1.0 NL/min), He carrier gas (~0.3–0.5 NL/min), and Ar sample gas (1.1–1.3 NL/min). Isotope data were collected in E-scan mode with magnet set at mass 202, and radio frequency power at 1245 W. Isotopes measured include 202Hg, 204(Pb+Hg), 206Pb, 207Pb, 208Pb, 232Th, and 238U. All isotopes were collected in counting mode with the exception of 232Th and 238U which were collected in analogue mode. Analyses were conducted in an ~40 min time resolved analysis mode. Each zircon analysis consisted of a 20 second integration with the laser firing on sample, and a 20 second delay to purge the previous sample and move to the next sample. Approximate depth of the ablation pit was ~20–30 microns.
The primary standard, 91500, was analyzed every 10–15 analyses to correct for in-run fractionation of Pb/U and Pb isotopes. Second zircon standards, Temora-2 and R33, were analyzed every ~10 analyses to assess reproducibility of the data. U-Pb analysis of Temora-2 and R33 during analytical sessions yielded concordant results and error-weighted average ages of 419.5 ± 3.6 Ma (n = 24) and 420.9 ± 1.6 Ma which is within close agreement of the accepted ages of 417–418 Ma and 417 Ma, respectively (Black et al., 2004; Mattinson, 2010). Corrections for minor amounts common Pb in zircon were made following methods of Tera and Wasserburg (1972) using measured 207Pb/206Pb and 238U/206Pb ratios and an age-appropriate Pb isotopic composition of Stacey and Kramers (1975). Zircons with large common Pb corrections (e.g., analyses interpreted as having ~10% or greater contribution from common Pb) were discarded from further consideration. Zircon dates are reported using the 206Pb/238U date for zircons <1400 Ma, and the 207Pb/206Pb date for zircons >1400 Ma. The quoted dates in the text and Table 1 are assigned 2% total uncertainties based on reproducibility of standards during analyses. For internal comparison of data from this study, 206Pb/238U dates shown on maps are assigned internal uncertainties only.
The results of U-Pb zircon analyses are summarized in Table 1 and geochemistry of intrusive rocks is summarized in Figure 6. Interpretations of this data are presented below under the appropriate road log stops.
ROAD LOG FOR PART 1
All coordinates in Part 1 are in WGS84 from Google Earth.
From Pasadena, follow Interstate 210 west and merge onto Interstate 5 North toward Santa Clarita and eventually to the Frazier Mountain Park Road exit, just north of Tejon Pass. Follow Frazier Mountain Park Road west to Lake of the Woods and turn left on Lockwood Valley Road. Turn left off of Lockwood Valley Road onto Frazier Mountain Road, where the Chuchupate Ranger Station is located on the right. Follow Frazier Mountain Road south where it turns into Route 8N04, continues past Chuchupate Campground, and switchbacks up Frazier Mountain. This section of road is graded dirt with some rocky sections and is passable for most two-wheel-drive vehicles with moderate clearance. (Note that this section of road is closed during the winter.) Stay right on Route 8N04 at the junction with Route 8N24, and continue south toward the edge of the mountain. Stay right to follow the road west along the ridge to a small pullout on the right at Stop 1.1a (34.7662, −118.9811).
Geologic Overview of Route to Stop 1.1
Interstate 210 skirts the southern flank of the western San Gabriel Mountains, which are underlain primarily by a complex array of crystalline basement rocks ranging from Paleoproterozoic to Cretaceous in age. The field trip will visit two study areas in this section of the range later in the afternoon of Day 1 (Placerita Canyon) and on the morning of Day 2 (Limerock Canyon). Just west of the merger of I-210 with I-5, the freeway crosses the mountain front fault zone at the transition from the San Fernando fault zone on the east to the Santa Susana fault on the west. The San Fernando fault was the source of the 1971 Mw 6.5 San Fernando earthquake that caused portions of the I-5/SR-14 interchange to collapse; the high overpass collapsed again during the 1994 Mw 6.7 Northridge earthquake. In the pass between the San Fernando and Santa Clarita Valleys, sedimentary rocks of the late Miocene to Pliocene marine Towsley and Pico Formations are exposed in the adjacent hills. The hills bordering Santa Clarita along I-5 are underlain primarily by Plio-Pleistocene, nonmarine, weakly indurated, sedimentary rock of the Saugus Formation and Pleistocene deposits of the Pacoima Formation. The large freshly graded pads and slopes visible west of the Magic Mountain theme park mark the first phase of development for the planned Newhall Ranch community.
South of Castaic Junction and the intersection with SR-126, the I-5 freeway crosses the projected trace of the Holser fault, and north of the junction I-5 crosses the San Gabriel fault, both of which are concealed below the alluvium of the Santa Clara River and Castaic Creek near their confluence. The San Gabriel fault is considered an early strand of the SAF system that accumulated 40–60 km of right-lateral separation between 12 and 5 Ma but has subsequently been reactivated in the Santa Clarita area during Pliocene and Quaternary transpression associated with the big bend in the SAF to the north. Castaic dam is visible upstream to the north on the east side of the town of Castaic. The California Department of Water Resources constructed this embankment dam as part of the California State Water Project; it opened in 1973, is 340 ft high, and has a total capacity of 325,000 acre-ft.
North of Castaic where the road grade starts to climb, rocks bordering the I-5 freeway transition from the Saugus Formation into underlying late Miocene strata of the marine Castaic Formation across an unconformable contact. The Castaic Formation marks the beginning of deposition within the late Miocene to Pliocene Ridge Basin, which encompasses all of the strata exposed along I-5 northward to the SAF. About 2 miles before the Templin Highway off ramp at Paradise Ranch, I-5 crosses the conformable contact between the Castaic Formation and the overlying nonmarine strata of the Ridge Basin Group. Here the I-5 roughly follows the axis of a macroscopic, northwest-plunging syncline that folds the entire Ridge Basin sequence, such that the view is down-structure and progressively younger strata are well exposed as you drive northward.
The nonmarine Ridge Basin Group is composed of three primary stratigraphic elements: Sandstone and conglomerate of the Ridge Route Formation on the northeast side of the basin, which interfinger to the southwest with a series of lacustrine members of the Peace Valley Formation near the axis of the basin, and the Violin Breccia, which forms an apron of debris fans along the northeast margin of the San Gabriel fault that interfinger to the northeast with the Ridge Route and Peace Valley Formations. I-5 transects interfingering members of the Ridge Route and Peace Valley Formations near the basin axis. The Violin Breccia is visible as dark-colored resistant outcrops on ridgelines to the southwest on the southwest limb of the syncline. The breccia is composed primarily of gneissic clasts derived from the Frazier and Alamo Mountain areas on the southwest side of the San Gabriel fault, which parallels the I-5 freeway ~1 mile to the southwest. Crowell (e.g., 1982) developed a conveyor-belt model of deposition along a progressively renewed scarp with a right-lateral, normal oblique sense of movement along the San Gabriel fault to explain the presence of ~14,000 m of strata along the southwestern margin of the basin based on a shingled depositional geometry such that the strata were not laid down in a single vertical pile and individual sediment packages pinch-out down-dip to the northwest against basement rock. The remainder of the sediments in Ridge Basin was derived primarily from sources to the north and northeast, and across the present-day Mojave strand of the SAF. May et al. (1993) developed an alternative model of basin formation and deposition assuming a listric geometry for the San Gabriel fault at depth below the basin.
Vista Del Lago Road provides access to a very informative visitor center and a beautiful overview of Pyramid Lake Reservoir, which is visible on the west side of I-5. This embankment dam reservoir, which was named after a pyramid-shaped rock outcrop created just south of the dam during construction of the Old Ridge Route, was completed by the California Department of Water Resources in 1973 as part of the California State Water Project; the dam is 386 ft high, and the reservoir has a capacity of 180,000 acre-ft.
Near the junction with Highway 138 (CA-138), an old quarry cut exploiting both Quaternary old alluvium and underlying granodiorite basement rock is visible on the west side of the highway. The large rounded peaks to the east are Bald and Liebre Mountains and are also underlain by this granodiorite but are commonly overprinted by landslide movement on their flanks. The granodiorite is part of the Cretaceous Liebre granodiorite exposed south of the SAF on the northeast margin of Ridge Basin. The road cut for Peace Valley Road just north of the quarry is the site of U-Pb sample WP-10L (see Table 1). Steeply north-dipping strata in this cut on the north side of the granodiorite represent the interfingering to gradational transition from the Ridge Route Formation to the overlying Lower Member of the latest Miocene to Pliocene–early Pleistocene(?) nonmarine Hungry Valley Formation. The prominent white interbed in this cut is composed of tephra; preliminary analyses indicate this tephra correlates with the 6.1 Ma Silver Peak Volcanics of Nevada (Elmira Wan, U.S. Geological Survey, 2018, personal commun.). Prominent outcrops of sandstone and conglomerate exposures with local greenish-gray mudstone interbeds exposed in the hills and ridgelines adjacent to the highway to the west from here represent both the Lower and Upper Members of the Hungry Valley Formation.
The prominent rounded ridgeline that fills the view to the north is Tejon Lookout ridge and lies on the north side of the SAF. This ridge is the type locality of the Cretaceous Tejon Lookout Granite, which is part of the southern end of the Sierra Nevada batholith; Sample WP-55L was collected from this ridge near Gorman Post Road for U-Pb dating (see Table 1); two other U-Pb samples were obtained from the southern flank of Tejon Lookout ridge farther to the west (WP-428 and WP-21L). Local purple-gray outcrops near the base of this ridge are andesitic volcanic rock slivers along the SAF that are tentatively correlated with the Oligocene or Miocene Neenach Volcanics sequence exposed farther to the east on the northeast side of the SAF. The Neenach Volcanics have been correlated with the Pinnacles Volcanic sequence to the north on the southwest side of the SAF to demonstrate a total right-lateral slip of ~314 km on the SAF since Oligocene time (e.g., Matthews, 1973; 1976). The SAF trends west-northwest here through the towns of Gorman and Frazier Park, contrary to the more typical northwesterly trend of the SAF north of the big bend in the Coast Ranges.
Farther northwest and south of I-5, a broad ridge descends toward I-5 just southeast of the town of Gorman. This ridge is composed of disrupted blocks of granodiorite, gneiss, marble, and sheared mafic metamorphic rock, which overlie the Lower Member of the Hungry Valley Formation. This block was interpreted as a thrust sheet flowering off of the SAF on older geologic maps (e.g., Crowell, 1982). More recent mapping by Weber (1988), Crowell (2002), and Olson and Swanson (2017) now recognize this as an ancient mega-landslide block that has subsequently been overprinted by late Quaternary slide movement back toward the north. Sample WP-91L of monzogranite/granodiorite was collected from a block in this slide for U-Pb dating (see Table 1). See description for Stop 1.4 below for additional details.
At Tejon Pass, the I-5 crosses over the main geomorphic trace of the SAF from the Pacific Plate to the North American Plate. The SAF is defined here by two primary strands that are separated by yellowish-brown weathering Quaternary sediments and juxtapose Tejon Lookout Granite on the north against the Hungry Valley Formation and small slivers and windows of light-gray weathered and pulverized Liebre Granodiorite to the south. (U-Pb sample WP-35L was obtained here from the Liebre Granodiorite.) Weathered granodioritic rock in cut exposures along Peace Valley Road locally appears to contain rounded granitic clasts of similar composition, which suggests this could represent an erosional contact with the overlying Hungry Valley Formation. Possible depositional relationships between the Hungry Valley Formation and underlying granodiorite were observed at two other locations south of the SAF and east of Gorman, which if correct, could locally define the northern depositional margin of the Hungry Valley Formation on basement rock. A sliver of Neenach(?) Volcanics is exposed in the road cut along the north side of the northern strand and rock exposed in the upper road cut east of the pass is largely overprinted by old slide movement.
Descending northwest from Tejon Pass, the view to the west is dominated by Frazier Mountain, which is underlain primarily by Paleoproterozoic biotite gneiss, quartzofeldspathic gneiss, and augen gneiss. Samples of splotchy banded gneiss (WP-527) and augen gneiss (WP-537) were collected from this flank of the mountain (see Table 1). The particularly prominent mesa surface on the mountain’s northeast flank is known as Condor Mesa and will be the site of Stop 1.4. The mesa is underlain by 50–75 m of pale-brown weathering Pleistocene fan deposits which are capped by a thin veneer of much younger fine-grained alluvium and underlain by pinkish-gray weathering sandstone of the Lower Member of the Hungry Valley Formation. Infrared stimulated luminescence (IRSL) and optically stimulated luminescence (OSL) dating is in progress to constrain the age of the Quaternary deposits mantling Condor Mesa. The northeast-facing slopes of Frazier Mountain northwest of Condor Mesa are underlain by ancient healed/mineralized rock avalanche landslide debris that locally underlies and pre-dates the Hungry Valley Formation, which will also be discussed at Stop 4.
The large valley north of Frazier Mountain is Cuddy Canyon and follows the SAF westward toward the towns of Frazier Park and Lake of the Woods. Elongate mesa surfaces bordering the south margin of the valley are underlain primarily by older alluvial deposits and are north of the main geomorphic trace of the SAF. The Frazier Mountain paleoseismic trench site explored by Lindvall et al. (2002) and Scharer et al. (2017) is located in a small extensional stepover along the SAF on the southwest margin of this older alluvium. The range on the north side of Cuddy Canyon is Tecuya Ridge, which is underlain the Cretaceous Lebec Granodiorite and associated schist, hornfels and marble pendant rocks. The I-5 bends easterly to the north to follow Castaic Valley, which conceals the main trace of the northeast-trending, left-lateral Garlock fault. This fault separates the Lebec Granodiorite from the Tejon Lookout Granite. Additional strands of the Garlock cut the Lebec Granodiorite north of the main trace, however, and these strands show more evidence of late Quaternary activity than the main strand. Sample WP-433 was collected from the Lebec Granodiorite between the main north and south strands of the Garlock fault for U-Pb dating.
After exiting I-5 onto Frazier Mountain Park Road, our route takes us past a low ridge that protrudes northward into Cuddy Canyon just east of the crossing of Cuddy Creek, a little over 2 miles west of the I-5. The surface of this ridge is at a similar elevation to the old alluvial surfaces mapped on trend to the southeast. However, detailed review of exposures in the channel cut-bank west of the crossing and other features reveal that this is ridge is composed of gneissic landslide debris derived from Frazier Mountain that has subsequently been offset by the SAF. This feature is the topic of Optional Stop 1.3 (see discussion below).
Opposite the intersection with Mt. Pinos Way, low hills at the base of Frazier Mountain are underlain by shattered Cretaceous granite, where sample WP-619 was obtained for U-Pb dating. The larger ridge to the west and south of Cuddy Creek opposite of Tecuya Mountain Road at Lake of the Woods is underlain by nonmarine strata and basalt correlated across the Lockwood Valley fault with the Plush Ranch Formation of Lockwood/Cuddy Valley; these exposures are separated from Frazier Mountain gneisses by the northeast-trending Cuddy Valley fault. The Plush Ranch basalts have been correlated with basalts in the Vasquez Formation of Soledad Basin and basalts in the Diligencia Formation of the Orocopia Mountains and used to estimate long-term displacement on the San Gabriel and San Andreas faults (e.g., Carman, 1964; Frizzell and Weigand, 1993). Farther to the west, before Lockwood Valley Road, Frazier Mountain Park Road crosses the concealed trace of the Lockwood Valley fault. This fault was previously interpreted as an extension of the left-lateral Big Pine fault (e.g., Hill and Dibblee, 1953), but Onderdonk et al. (2005) reinterpreted the segment in Lockwood Valley to be a separate fault.
Where Frazier Mountain Road reaches the northwest flank of Frazier Mountain, large blocks of augen gneiss are visible in the roadcuts, particularly at the tight hairpin turn. Carman (1964) interpreted these blocks to be part of the “Chuchupate” landslide, a Pleistocene mega-landslide that failed along the Cuddy Valley fault. Past the Chuchupate campground, the pavement ends and exposures of both augen gneiss and biotite gneiss of the Frazier Mountain Proterozoic gneisses are visible in cuts as the road switchbacks up the mountain side. The large mountain visible to the west across Lockwood Valley is Mt. Pinos, and the linear valley on its north flank follows the SAF toward Pine Mountain Club.
Stop 1.1a. Frazier Mountain: Paragneiss Sample Locality (34.7662, −118.9811)
On the south side of the road is a small outcrop of foliated and folded biotite paragneiss, the source of Sample WP-166B (see Table 1; Figs. 2 and 4). This sample yielded numerous zircons with a near unimodal distribution of U-Pb dates, generating an upper-intercept age of 1737 ± 35 Ma. The uniformity in individual zircon analyses and lack of older inherited zircons suggests the protolith may have been a tuff or other local silicic volcanic source unit within the parent sedimentary sequence of the paragneiss. This age corresponds well with previous 1750–1680 Ma age range estimated for Paleoproterozoic rocks in the western Transverse Ranges by Silver (1971), and is within the younger side of the error range age of the 1789 ± 61 Ma age reported by Barth et al. (2001) for the Mendenhall Gneiss northeast of the San Gabriel fault in the San Gabriel Mountains.
Hike 375 m (1250 ft) SSW down gentle terrain (~35 m downhill elevation loss) from Stop 1.1a to Stop 1.1b at the nose of a ridgeline overlooking Ridge Basin and the San Gabriel Mountains to the southeast.
Stop 1.1b. Frazier Mountain: Shear Zone in Augen Gneiss (34.7630, −118.9795)
At this site we will observe example outcrops of foliated augen gneiss that are overprinted by pervasive shear bands and mylonite, transforming the augen into flattened and dismembered porphyroclasts (Fig. 10A). The augen gneiss was sampled for U-Pb dating at WP-98B (site of Stop 1.2). The augen gneiss is interpreted to represent a protolith of porphyritic granodiorite or quartz diorite that intruded the surrounding paragneiss protolith at ca. 1700 Ma and was then metamorphosed later in the Proterozoic and overprinted by foliation that is developed at spatially variable intensity. The age of metamorphism and associated development of foliation is constrained by an unfoliated granite sampled at WP-97B (just south of Stop 1.2) that yielded a U-Pb crystallization age of 1431 ± 29 Ma (Fig. 7B, and see additional discussions for Stop 1.2 and the Results section of the guidebook above).
The distinct shear bands and mylonitic zones are oriented roughly parallel or slightly oblique to the metamorphic foliation here. The outcrops display a vague streaky lineation along the foliation that plunges 25°, N30E in a foliation plane that dips 38–46° to the east. Preliminary field assessment of S-C fabrics indicates top-to-the-southwest sense of movement. Thin sections from oriented samples are in preparation to further assess the slip kinematics of the shear bands exposed at this site and elsewhere on Frazier Mountain.
High-strain shear bands and mylonitic zones with streaky lineation were observed at numerous outcrops from the summit of Frazier mountain to the west, forming localized zones sub-parallel to the foliation with apparent increasing frequency and intensity toward the summit (Fig. 10B). Similar high strain rocks were also observed to the south near the base of the range in proximity to the Frazier Mountain fault. Observed streaky lineations consistently plunge to the northeast. Additional thin sections from oriented samples are in preparation to confirm the sense of slip and style of shearing.
The regional significance of this shear zone is uncertain at this time. However, gneiss observed on the southeast side of the mountain at least as far west as the Frazier Mine contains a significant component of splotchy-banded quartzofeldspathic gneiss (granulitic gneiss of Frizzell and Powell, 1982), a texture not observed west of the summit. The upper-intercept age of this distinctive, but currently undifferentiated subunit of Proterozoic gneiss found at WP-527 is 1634 ± 33 Ma, which is younger than the augen gneiss that intrudes the paragneiss west of the summit. The Xgn2 gneiss/schist sampled southeast of Frazier Mountain at WP-714B is also lithologically distinct and yielded an age of 1639 ± 33 Ma, also younger than the crystallization age of the augen gneiss. These relationships suggest there may be a significant lithologic, age, and/or metamorphic grade break caused by juxtaposition across a north-trending, west-dipping shear zone. Additional mapping is planned to elucidate the structural relationships near the summit of Frazier Mountain and additional U-Pb analyses are needed to further define the contrasting rocks on either side of the mountain.
The timing of activity on the shear zone is not directly constrained. However, the parallelism of the high strain bands with the metamorphic foliation suggests a possible temporal relationship. The metamorphic foliation is clearly truncated by the 1.43 Ga granite found at Sample site WP-97B. In addition, no high strain shear bands were observed within this small pod of granite, but neither are they apparent in the augen gneiss where it is cut by the granite. Hence, the absence of the shear bands in the granite does not prove they are pre−1.43 Ga age. Mapping by Zucker (1990) documented similar mylonite zones with top-to-the-north sense of slip along the Piru and Alamo shear zones cutting similar gneiss to the south of Frazier Mountain near Alamo Mountain (see Fig. 1). The eastern Piru shear zone roughly projects toward Frazier Mountain and the two zones may connect or be genetically related. Zucker concluded the Piru and Alamo shear zones were active in late Jurassic or early Cretaceous time based on (1) cross-cutting relationships with a younger Cretaceous pluton (Gold Hill), (2) the presence of undated but presumed Jurassic plutons (Stewart Mountain and Lockwood Creek) truncated by (or abutting against?) the shear zone, and (3) locally sheared granitic dikes of presumed Jurassic age within the shear zone. Interestingly, the mylonites mapped by Zucker are also generally parallel to the metamorphic foliation and form a wide zone within the gneiss, but are reportedly absent within the presumed Jurassic plutons. These relationships suggest that an older, pre-Jurassic period of mylonite formation may also be permissible.
More regionally, Miller and Wooden (1993) report the presence of a distributed zone of Proterozoic, north-trending, west-dipping mylonites in the New York Mountains with a top-to-the-east sense of shear that overprint 1660 Ma and older plutons and which become more pervasive near the crest of the range. Miller et al. (1986) note that an extensive mylonite zone separates the older metamorphic complex in the New York Mountains from relatively undeformed granitoid rocks to the east. It is hoped that additional mapping and analysis of the mylonite zones and presumed Jurassic plutons at Alamo Mountain will define their regional significance and confirm the age of shearing at Frazier Mountain and whether it is related spatially and temporally with the Piru and Alamo shear zones to the south.
Hike back to vehicles. If going to Stop 1.2, continue driving west to a point along the road at ~34.7631, −118.9891. Then hike down the southern slope of the mountain ~100 m in elevation on moderately steep terrain through patchy brush to 34.7601, −118.9877. Note that there is a section of road between Stop 1.1 and 1.2 that is steep and rutted and may be impassable for vehicles with low clearance or poor traction. If skipping Stop 1.2, drive back down Frazier Mountain and Lockwood Valley Roads to Frazier Mountain Park Road and return east either to optional Stop 1.3 (34.8188, −118.9291), ~2.5 miles west of I-5, or continue on to Stop 1.4.
Stop 1.2 (Optional). Frazier Mountain: Augen Gneiss Sample Locality; Intrusive 1.4 Ga Granite (34.7601, –118.9877)
This stop is at the location of foliated augen gneiss that is cut by alaskite sills and dikes that typically parallel the foliation. This is the site of U-Pb samples WP-98B (augen gneiss) and WP-98A (alaskite) (see Table 1). The augen gneiss sample yielded zircons with distinctive rims and cores. Analyses of the zircon cores produce an upper-intercept age of 1701 ± 34 Ma, interpreted as a crystallization age for the protolith. This result is very similar to the upper-intercept age of 1690 ± 5 Ma obtained in a 1997 analyses by Premo published in Kellogg (2001) for augen gneiss farther to the southwest on the Lockwood Valley Quadrangle. The augen gneiss age is also consistent with the age range of 1650–1680 Ma reported by Silver (1971) and within the error range of the upper-intercept age of 1679 ± 40 Ma reported by Barth et al. (1995a) for augen gneiss associated with the Mendenhall Gneiss northeast of the San Gabriel fault in the San Gabriel Mountains. Regionally, these dates also correspond to a period of voluminous pluton emplacement in the Mojave Desert region from 1.69 to 1.67 Ga, as noted by Wooden and Miller (1990). The Frazier Mountain augen gneiss is distinctly younger than the 1741 ± 35 Ma and 1731 ± 35 Ma ages obtained by Nourse (this paper, Part 2) for biotite granite augen gneiss exposed south of the San Gabriel fault at Limerock Canyon in the San Gabriel Mountains.
Analyses of the zircon rims in augen gneiss found in sample WP-98B result in an upper-intercept age of 1651 ± 33 Ma. This is slightly younger, but within the error range of the upper-intercept age of 1668 ± 33 Ma obtained for the alaskite intermixed along the foliation of the augen gneiss. Mapping to date suggests these alaskites occur both within the paragneiss and within the augen gneiss (Figs. 5B, 5C). The similar ages of the alaskite and zircon rims within the augen gneiss suggests the latter formed during high-temperature metamorphism that was either associated with intrusion of the alaskite and/or the metamorphism directly formed the alaskite by local partial melting of adjacent gneiss.
Silver (1971) suggested that the paragneiss was metamorphosed to at least amphibolite facies prior to intrusion of the augen gneiss protolith and then all of the rocks were deformed and metamorphosed to amphibolite and granulite facies at ca. 425–1450 Ma. However, only one generation of metamorphic foliation was observed within the augen gneiss and the surrounding paragneiss, and the foliation in both units appears to be roughly parallel overall. Therefore, the metamorphic event suggested by zircon rims and alaskite may also be responsible for producing foliation in the augen gneiss and for transforming the surrounding sediments into paragneiss. More regionally, the age of zircon rims and alaskite emplacement is coeval with a period of magmatism that occurred just prior to the Mazatzal orogeny at 1.67–1.69 Ga in the New Mexico area, as reported by Amato et al. (2008). Additional studies are needed to assess potential relationships between the Proterozoic terrane at Frazier Mountain and similar age terranes recognized to the east across the SAF in the southwestern United States.
A short distance south of the augen gneiss, a small body of fine- to medium-grained granite is exposed on a flat-topped ridgeline. This is the site of U-Pb sample WP-97B (see Table 1). U-Pb analyses resulted in an upper-intercept age of 1431 ± 29 Ma. This age correlates with anorogenic granites previously reported in the western United States (e.g., Anderson, 1983). The fine- to medium-grained texture, however, is not as coarse as typically reported elsewhere for this suite.
This granite shows evidence of a weak magmatic foliation but it is not overprinted by a metamorphic foliation. Relationships observed in the field clearly show that the granite cuts the metamorphic foliation observed in the adjacent augen gneiss (Fig. 7B). This places a minimum age of ca. 1.43 Ga on formation of the observed gneissic fabrics exposed on the southwest side of Frazier Mountain.
Stop 1.3 (Optional). Cuddy Canyon Holocene(?) Rock Avalanche (34.8188, −118.9291)
On the south side of Frazier Mountain Park Road, west of the bridge crossing Cuddy Creek, the creek has undercut the southern channel bank and exposed the geologic conditions below the typical colluvial cover. These exposures reveal highly disrupted, angular fragments of gneissic debris with textures indicative of a rock avalanche deposit, which will be discussed in more detail at the outcrop. These exposures extend across the mapped trace of the SAF, such that gneissic debris is present north of the fault trace. Previous mapping typically identified this material as older alluvium (e.g., Dibblee and Minch, 2006) or older alluvium overlying gneissic bedrock north of the SAF (e.g., Davis and Duebendorfer, 1987; Crowell et al., 2002). However, review of local exposures on the hill above the channel bank also reveal the presence of monolithologic angular fragments of gneiss and a lack of rounded clasts typical of alluvial deposits. The presence of gneiss north of the SAF here is problematic from a regional tectonic perspective.
Geomorphic evidence for past landslide movement at this location is suggested by a large scarp to the south on the lower flank of Frazier Mountain and an evacuated source area now occupied by young fan deposits and locally ponded sediments. We interpreted that a large rock avalanche failed across the SAF forming a lobe that protrudes into Cuddy Creek. This lobe has deflected the course of Cuddy Creek to the north, and the channel bank erosion at Stop 1.3 is in response to the associated bend in the stream. In addition, the landslide has subsequently been displaced right laterally by movement on the SAF (Fig. 11). The lateral margins of this slide provide piercing points across the landslide, which indicate right-lateral separation of ~200 m. Assuming an average slip rate of ~30 mm/year since it failed, the slide should be ~6700 years old. In order to provide an independent assessment of the slide’s age and the slip rate of the SAF since the landslide failed, samples have been collected from the surface of gneissic slide boulders exposed on top of the landslide for future cosmogenic exposure dating under the direction of Nic Barth.
Stop 1.4 is accessed by following Frazier Mountain Park Road eastward to Ralphs Ranch (Peace Valley) Road on the west side of I-5. Turn right and head south for ~2 miles to the entrance of the Hungry Valley State Recreational Vehicle Area at Gold Hill Road on the right. Head south to the entrance station and stop at the parking area to the right of the station to regroup and use the restroom facilities. (Note: the entrance station has maps that show and name the dirt roads accessing Condor Mesa.) Follow the dirt road west from the parking area (Stipa Trail) and follow it west past Badger Trail, and then veer left onto Condor Trail, which leads to the prominent elevated Condor Mesa. The low hills adjacent to the road are underlain by the Lower Member of the Hungry Valley Formation, the youngest formation of the Ridge Basin depositional sequence. The last section of road up to the mesa is moderately steep but is generally passable for two-wheel-drive vehicles with moderate clearance. Note that roadcuts below the mesa locally expose highly fractured splotchy banded gneiss. These rocks are disrupted by ancient (Pliocene?) landslide movement and were sourced from Frazier Mountain and either deposited onto the Lower Member of the Hungry Valley Formation or subsequently thrust over it along a northeast-trending, northwest-dipping thrust fault.
Stop 1.4. Condor Mesa Viewpoint and Lunch Break (34.7963, −118.9036)
From this spectacular viewpoint, we will have overviews of Frazier Mountain and associated Proterozoic gneiss bedrock to the southwest, the intersection of the San Andreas and Garlock faults to the north, Cretaceous Sierran-type plutons of the Lebec Granodiorite and Tejon Lookout Granite and associated pendant rocks exposed in the hills north of the SAF, and Cretaceous Liebre Granodiorite of the San Gabriel block exposed south of the SAF and east of Ridge Basin at Bald and Liebre Mountains to the east. New U-Pb zircon ages for these crystalline basement rocks and associated implications will be discussed at this stop. We will also review new mapping-based evidence for the distribution and age of large basement-derived landslides on the northeast flank of Frazier Mountain and along the south side of the SAF. Other prominent geologic features of interest surrounding this stop include the Pleistocene fan deposits underlying Condor Mesa, and the eastern strand of the Frazier Mountain fault, which lies at the base of the northeast flank of Frazier Mountain to the southwest of Condor Mesa.
Landslide Discussions—Northeast Flank of Frazier Mountain
Gneiss exposed on the northeast and southeast margins of Frazier Mountain are overprinted by multiple landslides ranging in age from pre–Hungry Valley time to Holocene time. At this stop we will discuss improved mapping-based understanding of three of the oldest slides on the south side of the SAF in more detail. These slides include two slide blocks mapped west of Condor Mesa, one on the northeast flank of Frazier Mountain and the second on ridgelines extending northeast toward Cuddy Valley (see map Fig. 3). The third is a very large slide block mapped to the east of Condor Mesa and southeast of Gorman (QTols on map Fig. 9). Two of these landslides were initially interpreted as shallow thrust plates splaying off of the SAF to the north (e.g., Crowell, 1982).
Old landslide debris mapped by Weber (1988) on the northeast flank Frazier Mountain exhibits a rubbly texture indicative of a rock avalanche failure mechanism (Figs. 12A, 12B). The debris is strongly overprinted by mineralization and is clearly cut by the east segment of the Frazier Mountain fault. These features suggest the slide is very old, and Weber interpreted the slide as initially failing prior to deposition of the Pliocene Hungry Valley Formation, a finding confirmed by Swanson and Olson (2016).
The ridge that extends northeast from Frazier Mountain toward the SAF is also underlain by slide-disrupted gneissic debris but is separated from the mountain front by outcrops of light-gray sandstone of the Hungry Valley Formation and older strata possibly correlating with the basal portion of the late Cretaceous to Paleocene San Francisquito Formation. This block was originally interpreted as a low-angle fault sliver flowering off of the SAF (e.g., Crowell, 1982), but has been reinterpreted as a landslide (Weber, 1988; Crowell et al., 2002; Swanson and Olson, 2016). New mapping by Swanson and Olson (2016) provides evidence that this is in fact a landslide block that pre-dates the Hungry Valley Formation but has subsequently been thrust back over the Hungry Valley Formation along a northeast-trending reverse fault. The initial phase of landslide movement likely occurred during uplift associated with late Miocene or early Pliocene activity on the San Gabriel fault. Conditions are further complicated by younger slide movement that locally overprints the older slide debris and the Hungry Valley Formation.
Three lines of evidence support the proposed interpretation that the landslide initially failed in pre–Hungry Valley time and the slide debris was later displaced by faulting:
A northeast-dipping contact exposed on the northeast flank of the slide block between steeply northeast-dipping layers of rubbly gneissic debris and the adjacent Hungry Valley Formation appears un-sheared and depositional in nature and cobble- to boulder-size clasts of weathered gneiss are present in the adjacent, concordant basal beds of the Hungry Valley Formation (see Fig. 12C). Younger strata of the Hungry Valley Formation contain few clasts of gneiss from Frazier Mountain. This contact is shown as a fault or landslide boundary on all previous published maps. These new observations substantiate that the landslide debris is older than the adjacent section of the Lower Member of the Hungry Valley Formation.
In apparent contradiction with this observation, the gneissic debris at the south end of the block clearly overlies the Hungry Valley Formation along a sheared contact (Fig. 13), which is shown as a landslide boundary on other published maps. If this contact was the result of landslide movement, then the slide would post-date the Hungry Valley Formation, in contradiction to the previous observation.
This apparent contradiction can be resolved by interpreting this contact as a younger fault. Evidence to support this conclusion begins with an exposure of a previously unmapped northeast-trending reverse fault to the southeast and associated northeast-trending overturned synclinal warps (photo in Fig. 14). Similar geologic structures exposed on either side of the canyon alluvium southeast of the slide suggest that a second parallel fault underlies this canyon (see map and cross section in Fig. 14). The southwest end of this concealed fault reasonably connects with the sheared contact exposed at the south end of the landslide block.
Landslide Discussions—Southeast of Gorman
The broad, northerly descending ridge visible to the east and south of I-5 and the town of Gorman is composed of disrupted blocks of granodiorite, gneiss, marble, and sheared mafic metamorphic rock, which overlie the Lower Member of the Hungry Valley Formation (Fig. 8). This block was interpreted as a shallow thrust sheet flowering off of the SAF on older geologic maps (e.g., Crowell et al., 1982). More recent mapping by Weber (1988) and Crowell et al. (2002), and confirmed by Olson and Swanson (2017), now recognize this as a large Plio-Pleistocene landslide block.
Based on the basement rock composition, the block could not have been sourced from the adjacent sedimentary rocks of the Hungry Valley Formation exposed south of the SAF, and the mixed rock assemblage is not compatible with the lithology of Frazier Mountain to the west, or the Tejon Lookout Granite and associated marble and hornfels pendant rocks currently exposed across the SAF to the north. However, the lithologies within this block are more compatible with clasts contained within the Upper Member of the Hungry Valley Formation, which include proximal cobble to boulder conglomerates, and Weber (1988) mapped local remnants of the Upper Member overlying the slide block. It is therefore inferred that this slide initially failed during deposition of the Upper Member of the Hungry Valley Formation and the source area to the north has now been offset across the Mojave segment of the SAF an indeterminate distance to the southeast.
No direct age is currently known for the Upper Member of the Hungry Valley Formation, but field evidence to the south demonstrates the presence of local angular discordance between locally steeply dipping beds of the Lower Member and more gently dipping beds of the overlapping Upper Member. The substantial change in clast content between the two members and the noted angular discordance suggest the Upper Member may be significantly younger than the Lower Member (upper Pliocene or early Pleistocene). A sample (WP-91L) of monzogranite/granodiorite exposed within the slide block was collected for U-Pb dating. Analyses yielded a crystallization age of 85.4 ± 1.7 Ma (see Table 1). This age is not diagnostic to discriminate between a source in the Sierra Nevada batholith or the San Gabriel block. However, analyses of five individual zircons yielded Proterozoic dates, indicating inheritance from Proterozoic country rock. This suggests the slide block is more closely associated with the Transverse Ranges than the Sierra Nevada batholith currently juxtaposed across the SAF to the north, and likely initially failed relatively early in the history of movement on the Mojave strand of the SAF.
Geomorphically young features such as north-facing scarps, backfilled grabens, and closed depressions are apparent within the northern portion of the slide block. These features indicate that subsequent to initial failure southward across the SAF, younger, northward-failing landslide movement has overprinted the older landslide block. Additional closed depressions with ponded sediment are also present on the southern portion of the slide mass. The origin of these depressions is enigmatic as they are relatively youthful and not likely related to the original landslide failure as such features would not likely have been preserved during deposition of the Upper Member. Likewise, there is not clear independent evidence that the southern portion of the old slide block has moved back toward the north.
Overview of Cretaceous Plutons
From Condor Mesa there is an expansive view to the north across the SAF of Tecuya Ridge, which is underlain by the Cretaceous Lebec Granodiorite and associated pendant rocks of marble, schist, and hornfels, and of Tejon Lookout Ridge across the I-5 freeway to the northeast, which is underlain by the Tejon Lookout Granite and its associated pendant rocks of marble and hornfels. The plutons are considered part of the southern Sierra Nevada batholith. The rounded mountains across Ridge Basin to the east on the south side of the SAF are Liebre and Bald Mountains, which are underlain in part by the Liebre Granodiorite. Unnamed Cretaceous granites are also exposed near the northwest and southeast flanks of Frazier Mountain (Kg1 and Kg2 respectively). The photo in Figure 4B illustrates Kg2 granite cutting biotite gneiss/schist at WP-714 southeast of Frazier Mountain. The plutons south of the SAF are in the western San Gabriel Mountains block of the central Transverse Ranges.
Samples were collected from each of the major Cretaceous bodies for U-Pb zircon dating and geochemical analyses (see map Figs. 3, 8, and 9 maps). U-Pb data are summarized in Table 1 and the geochemical data are summarized in the TAS plot in Figure 6. One goal of this work was to advance the understanding of the age and geochemical character of the subject rocks for purposes of basic geologic mapping. A second goal was to assess if the plutons on either side of the SAF and Garlock fault have distinct ages or geochemical signatures. The affinities of isolated outcrops and fault slivers of granitoids exposed along the SAF have been unclear, which hinders recognition of the main strand of the SAF with large-scale, long-term slip.
In this respect, the analyses were successful in identifying isolated outcrops near Tejon Pass as properly assigned to the Liebre Granodiorite. Analysis of WP-35L, which was previously assigned to the Tejon Lookout Granite, matched the age and petrography of the Liebre Granodiorite, but did not show Proterozoic inheritance signature. Individual zircon analyses demonstrate that the Kg1 and Kg2 granites on the flanks of Frazier Mountain display a strong inheritance signature of older rocks extending into the Paleoproterozoic, which is not surprising based on the proximity of the Frazier Mountain gneiss country rock. WP-10L from the Liebre Granodiorite shows a strong Proterozoic inheritance signature as expected. Samples from north of the SAF (WP-428, WP-433, WP-21L, and WP-55L), did not show an inherited Proterozoic signature. Sample WP-61L located along the southern margin of the SAF yielded a few individual grains with younger Proterozoic inheritance signature.
Pleistocene Alluvium Underlying Condor Mesa
Stop 1.4 is on Condor Mesa, the northeastern face of which exposes a 50- to 60-m-thick section of Pleistocene alluvium containing abundant gravel to small boulder-size, angular clasts of gneiss derived from Frazier Mountain. The coarse alluvium overlies pinkish-gray sandstone of the Hungry Valley Formation and is capped by a thin veneer of fine-grained, poorly consolidated alluvium with poor soil development. A soil stratigraphy study by Zhao (1989, 1990) noted anomalously poor soil development on this surface with respect to the elevated position of the mesa and extensive dissection but did not identify the younger capping unit. New mapping by Swanson and Olson (2016) recognized the presence of this younger capping unit, which is sourced from a canyon to the south. Luminescence dating is in progress to better understand the ages of the thick Pleistocene deposits and the younger capping alluvium.
East Frazier Mountain Fault—The Offset, Reactivated Continuation of the San Gabriel Fault?
Current regional geologic maps show Frazier Mountain as bounded to the south and east by the Frazier Mountain fault system, which juxtaposes gneissic rock of Frazier Mountain against more gneiss and the Pliocene Hungry Valley Formation. The inferred low-angle, thrust fault geometry for the Frazier Mountain thrust fault dates to the original interpretation by Buwalda et al. (1930). The low-angle lobe mapped by Buwalda is now interpreted as a landslide deposit, however, and a true outcrop of the southern fault plane could not be found during recent mapping to confirm the dip of the fault due to extensive overprinting by landslide movement, creep, and colluvial drape. The map trace bends 90° from east-northeast–trending to north-northwest–trending at the southeastern corner of the mountain (see Fig. 1) and the northwest-trending segment of this fault, referred to herein as the East Frazier Mountain fault, is visible from Stop 1.4 at the base of Frazier Mountain southwest of Condor Mesa. In more detail, field observations support that the East Frazier Mountain fault dips to the northeast away from Frazier Mountain rather than under it. Crowell (2003a) interprets the East Frazier Mountain fault segment as an exhumed and reactivated segment of the San Gabriel fault offset roughly 2 km and rotated to the east by the Frazier Mountain thrust fault rather than a bent manifestation of the thrust fault (e.g., Crowell, 1950; Crowell, 2003a, cross section 14). Crowell (1982; 2003a) interprets that the Miocene San Gabriel fault was reactivated in the Pliocene during renewed uplift of Frazier Mountain, which is supported by the presence of local beds containing abundant clasts of Frazier Mountain gneiss within the Hungry Valley Formation adjacent to the fault, similar in character to the older Violin Breccia.
At a regional scale, the San Gabriel fault is thought to have accommodated 40–60 km of right-lateral oblique slip acting as the primary plate boundary from ca. 12–5 Ma. The path of the Miocene-age San Gabriel fault trace, however, is obscured by later faulting and deposition on the east margin of Frazier Mountain as it approaches the SAF. Crowell interpreted that the Miocene manifestation of the San Gabriel fault is concealed at depth below Frazier Mountain by the Frazier Mountain fault (and Dry Canyon thrust to the south), which provides a path for the San Gabriel fault to join with the SAF, as required by regional plate boundary evolution models. Cross section 14 in Crowell (2003a) correctly shows the East Frazier Mountain fault dipping to the northeast at the surface (as verified in the field) and juxtaposing the Hungry Valley Formation against Frazier Mountain gneiss. In detail, however, his cross section interprets that the low-angle Frazier Mountain thrust fault connects to the northeast-dipping eastern fault at depth, but that the thrust does not extend up section past the San Gabriel fault strand to the northeast into the Hungry Valley Formation. The lack of this extension beyond the gneiss at depth or of a trace mapped at the surface extending to the east creates concerns regarding the structural viability of this model as the 2 km of reverse slip shown in the footwall of the east Frazier Mountain fault would have to be accommodated by folding or bulk shortening up dip in the adjacent Hungry Valley Formation. In addition, the resulting compression would be expected to produce northeast-side-up reverse movement on the northeast-dipping East Frazier Mountain fault rather than the observed normal separation. Updated mapping (Swanson and Olson, 2016) shows that structure in the Hungry Valley Formation northeast of the East Frazier Mountain fault is dominantly perpendicular to the fault with subordinate structures parallel to it. The structural geometry of the Hungry Valley Formation east of the interpreted San Gabriel fault strand does not appear to validate 2 km of shortening. Weber (1986, 1988) questioned whether the eastern strand is a reactivated segment of the San Gabriel fault. Further study is needed to unravel the Pliocene histories and interactions of the San Gabriel and Frazier Mountains faults.
Crowell (2003c) interprets the Frazier Mountain fault as initiating in Pliocene time during deposition of younger strata of the Hungry Valley Formation. However, a series of newly mapped outcrops of late Oligocene rhyolitic volcanic diatreme rocks and tuff (preliminary Argon date of 25 Ma) coincide with the fault trace on the south side of Frazier Mountain. This alignment strongly suggests that a fault was present at this location during the early Miocene, long before development of the San Gabriel fault. The volcanic age is consistent with previously reported dates of the Plush Ranch and Vasquez Formation basalts (Frizzell and Weigand, 1993), which erupted during a time of widespread Miocene transtension in the Transverse Ranges. This suggests the Plio-Pleistocene phase of movement on the Frazier Mountain fault may have involved reactivation of a Miocene, low-angle normal fault that was the locus of small volcanic vents at the time. Implications of the rhyolitic composition at this time and presence of a Miocene extensional fault at Frazier Mountain are topics of ongoing research.
Return to vehicles, and drive back to Interstate 5. Drive south to the junction with the 14 Freeway to begin road log for Part 2 of this field trip.
PART 2. GEOLOGY AND U-Pb ZIRCON GEOCHRONOLOGY OF THE PLACERITA FORMATION AND ASSOCIATED INTRUSIVE ROCKS, WESTERN SAN GABRIEL MOUNTAINS (DAY 1, P.M.; DAY 2, EARLY A.M.)
LOCATION AND SIGNIFICANCE
The Placerita Formation in the western San Gabriel Mountains is a distinctive metamorphic assemblage of presumed Paleozoic or Precambrian age situated at the southwest margin of the Laurentia in tectonic reconstructions. Early studies by Miller (1934) and Oakeshott (1958) described a northeast-dipping, 800-m-thick metasedimentary package intruded by Paleozoic(?) and Mesozoic(?) plutons. The metamorphic assemblage includes upper amphibolite facies pelitic gneiss, quartzite, calcsilicate gneiss, and marble that exhibit folding on regional and outcrop scales. Individual lithologic units were not subdivided at the reconnaissance scale of mapping, nor have these rocks been directly dated.
We conducted a detailed geological mapping and geochronology study of the Placerita Formation in Placerita Canyon and Limerock Canyon (Fig. 15; Vermillion and Nourse, 2018, 2019; Murphy et al., 2018; Nourse et al., 2019). These localities were denuded by the July 2016 Sand Canyon and December 2017 Creek wildfires, respectively, allowing excellent exposure of field relationships previously covered by dense brush. Below we present geologic maps of the two study sites, in addition to descriptions of the stratigraphic units and structural geometry. We also report new U-Pb zircon ages acquired on the sensitive high-resolution ion microprobe–reverse geometry (SHRIMP-RG) and laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) that constrain key age relationships. We demonstrate that the sedimentary protoliths were likely deposited in the Neoproterozoic, with provenance dominantly sourced from southwestern Laurentia. Interlayered, variably foliated, and cross-cutting intrusive rocks are shown to be Early Jurassic, Late Jurassic, and Late Cretaceous, respectively. We demonstrate a Late Jurassic age for the principle upper amphibolite grade metamorphic/deformation event.
Geologic relations and U-Pb geochronology presented in this field trip guidebook lay the framework for future publications on paleotectonic reconstructions of southern California. For example, the Placerita Formation provides important piercing points across various strands of the San Gabriel and San Andreas fault systems. In a pre–San Andreas restoration, these Neoproterozoic continental margin metasediments occupy a position at the far southwest edge of Laurentia that was undergoing active rifting. Detrital zircons within this formation offer new provenance constraints on various models for the configuration and break-up of the Rodinia supercontinent during the Neoproterozoic (see Whitmeyer and Karlstrom, 2007, and references therein). Such discussions, along with detailed presentation of supporting geochronological data, are beyond the scope of this article.
The western San Gabriel Mountains preserve a distinct marine metasedimentary assemblage (Placerita Formation) that forms part of a framework for younger foliated and unfoliated intrusive rocks. Previous works on the metamorphic rocks yielded abundant descriptions of lithology and metamorphic mineral assemblages, but significant uncertainty remains about the age of the sedimentary protoliths and the timing of plutonism and metamorphism. Likewise, none of the published studies provide much insight into the geometry of folds that are abundant and clearly record significant tectonism. Below we summarize the preexisting state of knowledge that is in some cases conflicting.
Miller (1934) first described the metasedimentary rocks observed in a 1:100,000-scale mapping endeavor and proposed the name “Placerita formation” for occurrences in Placerita Canyon and Los Pinetos Canyon (Fig. 16). There he documented a steeply northeast-dipping section of “biotite schist, quartzite, quartz-pyroxene gneiss, and crystalline limestone,” commonly “injected by granite and dioritic material.” The Placerita Formation in this type area has a minimum thickness of 2500 ft. Miller also described lenses of similar rock farther southeast in Little Tujunga Canyon, and in the Verdugo Mountains and San Rafael Hills, where sillimanite schist was observed (Fig. 15). Several occurrences of graphitic schist are also noted in these areas.
Despite the absence of fossils in the metasedimentary strata and lack of modern geochronology, Miller (1934) made some astute observations pertaining to age constraints. He described field relationships (p. 34–35) where xenoliths of sillimanite schist are included in unfoliated “Wilson diorite” of presumed Late Jurassic–Early Cretaceous age. In several places, sill-like bodies of foliated diorite and granite (inferred to be Paleozoic age) are interlayered with Placerita quartzite and schist but sharply intruded by Wilson diorite (p. 53–54), implying that strong metamorphism represented by the foliated units occurred during late Paleozoic or early Mesozoic time. Miller also speculated that the Placerita Formation with its “lit-par-lit” injections of granite is older than Pelona schist of nearby Sierra Pelona (Fig. 15). He preferred a Precambrian depositional age, distinguishing the Placerita Formation from marine sedimentary strata designated by others as Paleozoic in the San Bernardino and San Jacinto Mountains (Miller, 1934, p. 63–64).
In his geologic map of the 15 min San Fernando quadrangle, Oakeshott (1958) denoted multiple occurrences of Placerita Formation that correspond to areas previously mapped by Miller (1934), along with a few foliation measurements indicating a general northwest strike. He provided additional petrographic details in the corresponding report and emphasized the presence of graphite in many of the schists and limestones. Oakeshott indicated a “Triassic or older” age for the metasediments based on “intrusion of the series by Cretaceous(?) and Jurassic(?) granitic rocks and older dioritic gneiss.” He suggested correlation of the many limestone lenses (distinguished as Pl on his map) with the Mississippian Furnace Creek Limestone of the San Bernardino Mountains; however, he also noted that “an earlier Paleozoic or Precambrian age for the Placerita Formation cannot be precluded” (Oakeshott, 1958, p. 52).
Ehlig (1975b) and Dibblee (1982b) recognized the Placerita Formation as an important unit in the San Gabriel Mountains block. Their geologic descriptions mimicked those of Miller (1934) and Oakeshott (1958), but an important new contribution was the extension of likely correlative rocks to areas of the southeastern San Gabriel Mountains, specifically Potato Mountain and Ontario Ridge (Fig. 15). Ehlig’s 1975 map, used as the base for Figure 15, shows occurrences of Placerita-like metasedimentary strata as a hybrid unit: “Mesozoic granitic rocks with Paleozoic(?) metasediments and gneiss.”
A more detailed description of the western Ontario Ridge metasedimentary sequence is provided by May (1986) and May and Walker (1989), who referred to these rocks as “San Antonio terrane.” Similar rocks on eastern Ontario Ridge and Penstock Ridge (Fig. 15) are mapped at 1:24,000 scale on U.S. Geological Survey geologic maps of Cucamonga Peak quadrangle (Morton and Matti, 1990a) and Devore quadrangle (Morton and Matti, 1990b). These authors lump the metasedimentary strata into one unit (“Schist and Gneiss [Paleozoic]”).
Powell (1993) proposed a formal regional name “Limerock Canyon assemblage” for rocks similar to the metasedimentary and metaigneous succession of Limerock Canyon (where the best exposures of Placerita Formation occur, according to Oakeshott, 1958). He suggested correlation to multiple bodies of nonfossiliferous marine metasedimentary rocks presently separated by strands of the late Cenozoic San Gabriel and San Andreas fault systems (see p. 82 of Powell (1993) for a complete list of localities). Powell used the Limerock Canyon assemblage along with multiple distinctive Proterozoic and Mesozoic basement rocks to develop a palinspastic reconstruction of the San Andreas system in southern California. In its type area at Limerock Canyon, the section is composed of “graphite schist, vitreous quartzite, pelitic schist, marble, and calcsilicate rocks typically intruded by and intermingled with granitic orthogneiss and amphibolite” (Powell, 1993). An upper-amphibolite facies metamorphic mineral assemblage is “characterized by sillimanite-garnet-cordierite and cordierite-orthoamphibole-bearing pelitic rocks, plagioclase-diopside-garnet ± scapolite and plagioclase-wollastonite-scapolite–bearing calcsilicate rocks, and diopside-forsterite, diopside-wollastonite, and scapolite-bearing calcite marble.”
Age relationships involving the Limerock Canyon assemblage are summarized from Powell (1993, p. 82) as follows: “The age of the protolith is demonstrably older than the Late Cretaceous plutons that intruded it between about 75 and 90 Ma. Relations with Triassic and Proterozoic plutonic rocks in the San Gabriel and San Bernardino Mountains are obscured by the intrusion of Cretaceous plutonic rocks and the accompanying deformation and by late Cenozoic strike-slip faulting. It seems most likely that the protolith of the Limerock Canyon metasedimentary suite is Paleozoic or Proterozoic in age, and there is evidence for either interpretation.” (Powell goes on to list arguments for either age assignment. Regarding the timing of metamorphism, he states, “Mineral paragenesis and microfabrics of the Limerock Canyon metamorphic rocks appear to have formed in a single metamorphic event, which is reasonably inferred to be related to Cretaceous plutonism and tectonism” (Powell, 1993, p. 82).
Finally, a recent geochronology study of quartzites from the Ontario Ridge metasedimentary sequence, and associated intrusive rocks, is detailed in the M.S. thesis of Zylstra (2017). Results of this thesis suggest a Neoproterozoic depositional age for the sedimentary strata and provide important context for comparison to the type Placerita Formation described below.
PLACERITA AND LOS PINETOS CANYONS
Figure 16 presents our new geological mapping from the type area of the Placerita Formation, especially well exposed in Placerita Canyon and Los Pinetos Canyon.
Geologic Map Relations
East-trending Placerita Canyon cuts obliquely across the type section of Placerita Formation described by Miller (1934). Three north and northwest-trending tributaries of Los Pinetos Canyon (described briefly by Miller, 1934, and Oakeshott, 1958) cut the section at high angles, providing additional cross-strike control. These canyons and several of the ridge lines were walkable during our mapping endeavors of 2017–2019. The geologic map of Figure 16 shows the distribution and geometry of mapped rock units. Our map bears some resemblance to the larger scale reconnaissance geologic maps produced by Miller and Oakeshott, but significant differences exist, particularly in delineation of the younger intrusive rocks.
In general, metasedimentary components of the Placerita Formation strike west-northwest and dip ~60°–75° northeast. Multiple dip reversals indicate folding along northwest trends. Younger intrusive bodies, both strongly foliated and weakly foliated, tend to be tabular and emplaced parallel to layering in the metasedimentary host rocks. In the southeast headwaters of Los Pinetos Canyon, a large, irregular-shaped body of unfoliated biotite granite intrudes the east end of the exposed Placerita Formation. Farther west, and south of Placerita Canyon, this granite is intruded by a large body of weakly foliated biotite hornblende quartz diorite. The quartz diorite forms the south margin of the main Placerita Formation exposures, but several elongate xenoliths of metasedimentary strata occur within its interior. The quartz diorite also contains abundant elongate xenoliths of fine-grained, recrystallized dioritic gneiss, and less common xenoliths of mylonitized biotite granite.
Within the main part of Placerita Canyon, two sills of weakly foliated quartz diorite similar to the larger body farther south intrude the Placerita Formation metasedimentary units. Also in this area is a sill-like mass of foliated biotite granite with xenoliths of biotite gneiss.
The rugged area between lower Placerita Canyon in the west and Lost Pinetos Canyon on the east was not mapped due to thick brush cover spared by the recent Sand Canyon wildfire. Google Earth satellite imagery, however, allowed projection of some quartzite and granite marker units between the two areas.
It is tempting to treat the Placerita Formation as an intact stratigraphic section, but existence of many tight to isoclinal folds at outcrop scale, and several map-scale dip reversals lead one to be cautious. Parts of this section are likely duplicated. It is also difficult to ascertain whether the section is upright or overturned. Nevertheless, the general stratigraphic pattern revealed by the geologic map suggests that a quartzite-rich interval with some marble and biotite gneiss is overlain by a thick calcsilicate unit with marble, in turn overlain by a section dominated by biotite gneiss with subordinate quartzite and marble. In terms of simple sedimentary protoliths, this might represent a prograding shoreline sequence with sands overlain by calcareous muds and limestone, overlain by clay-rich shales and siltstones.
The Placerita metasedimentary and plutonic complex is overlapped on the west, north, and northwest by Neogene sandstone and conglomerate (Battle, 1939; Oakeshott, 1958). Numerous Quaternary landslides and colluvial deposits cover many of the ridges south of Placerita Canyon, obscuring the connectivity between some map units that are projected between tributary canyon exposures.
Presented below are field relationships and petrologic descriptions of the rock units distinguished on Figure 16. A limited selection of thin sections provides additional information on the metamorphic mineral assemblages and textures. Selected photographs (Figs. 17A–17I) give additional field context.
Metasedimentary Rocks of the Placerita Formation
Biotite Gneiss (Map Unit bgn)
The most abundant lithology in the Placerita Formation is a biotite-quartz-feldspar gneiss with distinct banding that reflects alternation of biotite-rich and quartzofeldspathic layers. Typical banding occurs on a scale of 5 mm to a few centimeters although some biotite-rich layers approach a meter in thickness (Figs. 17A and 17B). Texture varies from gneissose to schistose, usually medium grained, but finer grained phyllite layers are common. Also present are layers of fine- to medium-grained foliated hornblende-biotite diorite or amphibolite. These represent intrusions, variably referred to as “Ruby diorite” by Miller (1934), “diorite gneiss” by Oakeshott (1958), or “amphibolite” by Powell (1993), that share foliation with the Placerita Formation host rocks.
The biotite gneiss occasionally displays visible almandine(?) garnet and sillimanite, confirming a pelitic composition affected by an upper amphibolite grade metamorphism. These minerals may escape detection by the casual observer but are present in some of the thin sections. Outcrop-scale folds are common, and a few dip reversals define map-scale folds (Fig. 16); see also structure section below.
Marble (Map Unit m)
White, coarsely crystalline marble is the least abundant metasedimentary unit but is most conspicuous in its outcrop appearance. This rock forms either continuous 1-m- to 5-m-thick layers interstratified with quartzite and calcsilicate or strings of resistant lenses within biotite gneiss or schist higher in the section. Thinner centimeter- to decimeter-scale layers of marble are common within the calcsilicate gneiss unit, where recessively weathered bands of calcareous rock alternate with more resistant calcsilicate or quartz-rich layers. Near the top of the exposed section, a distinctive marble marker, outcropping for ~3 km as isolated 1-m- to 10-m-wide pods, follows a resistant stratigraphic horizon surrounded by poorly exposed biotite gneiss and schist.
Oakeshott (1958) distinguished many outcrops of this unit as “white crystalline limestone” (Pls) on his map, either as isolated lenses within a mixed unit of Placerita Formation and dioritic gneiss, or as xenoliths in granodiorite. Petrographic descriptions by Oakeshott (see, also Powell, 1993) indicate that the dominant mineral phase is calcite, with lesser quantities of diopside, wollastonite, forsterite, and scapolite in the less pure varieties. Oakeshott also cites MgO measurements that indicate some of this unit is dolomitic. We cut only one thin section from a distinct spotted outcrop of white marble containing abundant 5–10 mm resistant dark garnets (sample JN 1830). The garnets are dispersed within a foliated matrix of calcite, epidote, plagioclase, and titanite.
Miller (1934) described additional isolated outcrops of marble farther east in the San Gabriel Mountain that are generally associated with quartzite and pelitic schist. These occurrences offer one basis for regional correlation of the Placerita Formation.
Calcsilicate Gneiss (Map Unit cs)
This unit corresponds to the quartz-pyroxene gneiss of Miller (1934). The thickest exposures that occur in the middle part of the stratigraphic sequence in Los Pinetos canyon are characterized by distinctive calcsilicate mineral assemblages and spectacular folds (Fig. 17C). Alternating green and orange bands rich in epidote and grossular garnet are common (Fig. 17D), as are light-colored hornblende plagioclase gneisses. Frequently observed in this unit are recessively weathered bands of calcite or dolomite interlayered on a centimeter scale with resistant layers of quartzite or calcsilicate minerals. The contrasting competence of these layers is conducive to fold development.
One thin section cut from a typical banded calcsilicate gneiss (sample JN 1829) revealed distinct layers rich in epidote, grossular, and hornblende-plagioclase, in addition to smaller quantities of quartz, calcite, and titanite. Calcsilicate gneisses are also associated in general with white marble bodies at deeper and shallower structural levels of the Placerita Formation. Only the larger occurrences are distinguished on Figure 16.
Quartzite (Map Unit q)
We mapped two varieties of quartzite or quartz-rich gneiss as resistant 20-cm- to 5-m-wide layers within the bgn and cs units. Zones of metasedimentary rock rich in quartzite are lumped together and distinguished as mappable quartzite layers on Figure 16. Zircon yield from four samples of quartzite was prolific, demonstrating that they originated as detrital sandstones rather than cherts. The thicker meter-scale bodies of quartzite (characterized by geochronology samples JN 1800 and JN 1721; Fig. 17E) are light to medium gray and vitreous where freshly broken. These typically occur within red- and orange-weathered zones of biotite gneiss interlayered with thin cm-scale quartzite bands. Thin sections revealed a compositional range from orthoquartzite (93% quartz) to meta-arkose with thin beds containing 30% to 60% feldspar. Coarse quartz grains form elongate domains, statically annealed, that define a weak foliation with aligned biotite. Traces of garnet are present along with the zircon.
A second variety of impure quartzite, represented by geochronology samples JN 1720 (Fig. 17F) and JNKV 1809, is intimately associated with calcsilicate gneisses. These quartzites form layers 1 cm to 30 cm thick, commonly folded, that stand out from adjacent calcareous strata. Thin-section examination shows these rocks to be titanite-anorthite-dolomite-diopside-grossular-epidote quartz gneisses with a statically recrystallized hornfelsic texture. Quartz content varies from 60% to 75%. JN 1720 contains wollastonite, which along with diopside, confirms metamorphic grade to be upper amphibolite facies. The bulk composition and abundant zircon suggests a calcareous sandstone protolith.
Mesozoic Intrusive Rocks
Several varieties of plutonic rocks, both foliated and unfoliated, intrude the Placerita Formation and provide constraints on the timing of metamorphism and folding. These rocks were generally considered to be Mesozoic by Miller (1934) and Oakeshott (1958) although the latter author left open possibility of a late Paleozoic age for the older dioritic gneiss.
Jurassic(?) Diorite (Map Units Jdi and Jdign)
This unit corresponds to the “Rubio Diorite” of Miller (1934) and the “diorite gneiss” of Oakeshott (1958), who offered the following description of this unit. “The rocks mapped as diorite gneiss comprise a variety of dark gneisses, metadiorites, massive hornblende diorite, and amphibole and biotitic schists. Some are clearly coarse- to fine-grained intrusives cutting the Placerita Formation metasedimentary rocks and injected into them; others may actually be part of the Placerita.” We recognize all of these rock types in the study area and concur that they represent the earliest generation of intrusive rocks.
In Los Pinetos Canyon, most of the dioritic rocks associated with the Placerita metasediments occur as 1–30 cm concordant layers of fine- to very fine-grained hornblende-biotite diorite or amphibolite, not mappable at the scale of Figure 16. These invariably display a weak to moderate foliation parallel to that in the host rocks, and in several places are folded with the metasedimentary strata, indicating that metamorphism and folding of the Placerita Formation post-dated intrusion of the diorite. A few larger irregular bodies of fine-grained hornblende diorite or amphibolite are mapped in the main branch of lower Los Pinetos Canyon. Although somewhat less foliated than their host rocks, these exhibit “pinch and swell” or incipient boudinage structures.
Elsewhere, the dioritic rocks record important field relations with younger plutonic rocks. A small body of weakly foliated, medium- to fine-grained hornblende diorite is intruded by Jurassic granite on the ridge separating the central and east branches of Los Pinetos Canyon. Farther southwest along Santa Clara Truck Trail, elongate xenoliths of well-foliated fine-grained hornblende-biotite diorite gneiss are abundant within the Late Cretaceous quartz diorite pluton that dominates the south-central part of the map area (Fig. 17G). The largest of these xenoliths is intruded by multiple sheets of biotite granite (Fig. 17H) that exhibit strong mylonitic fabric parallel to the diorite foliation. Sample JNKV 1807 was collected from one of these granite sheets.
Late Jurassic Biotite Granite (Map Units Jgr, Jgrgn, and Jhbqm)
A large body of biotite granite intrudes the eastern portions of the Placerita Formation. This pluton, represented by sample JNKV 1810, is generally medium grained and slightly porphyritic, and is only subtly foliated except near its contacts with metasedimentary host rocks. Composition ranges from biotite monzogranite to alkali granite, with an unmapped transition to hornblende-biotite quartz monzonite in the area southeast of May Canyon Saddle (Fig. 16). A coarse porphyritic phase of the alkali granite also exists (based on float block sample JN 1817), but the source exposure was not located. The Late Cretaceous Los Pinetos quartz diorite (described below) intrudes the granite and separates it into two parts. Elongate inclusions of granite are common in the quartz diorite.
We mapped several sills of foliated granite within the Placerita Formation farther west in the main part of Placerita Canyon. One of these (geochronology sample JN 1722, mapped as Jgrgn) displays a strongly lineated mylonitic fabric and contains inclusions of biotite gneiss. Additional occurrences of mylonitic granite gneiss were observed within the diorite gneiss exposure in the southwest part of the map area (Fig. 17H). Representative sample JNKV 1807 shares a common foliation with its diorite gneiss host and displays a cryptic southwest-vergent shear sense.
Six thin sections cut from various phases of the granite all preserve a pervasive thermal overprint manifested by recrystallization of quartz and biotite. Alkali feldspar is also recrystallized into subgrains in two of the mylonite samples. Mylonitic textures observed in hand specimen are statically annealed such that shear sense could not be confirmed. Even the very weakly foliated sample JNKV 1810 displays partially recrystallized quartz grains. Presumably these thermal effects are related to intrusion of the Los Pinetos quartz diorite.
The Late Jurassic granite provides important time constraints on metamorphism and deformation in the study area. Near sample site JNKV 1810, the very weakly foliated granite contains a 1200 m long screen of biotite gneiss that displays strong foliation and folding characteristic of the Placerita series. Elsewhere, two tabular bodies of granite (samples JN 1722 and JNKV 1807) that intrude the Placerita Formation and diorite units, respectively, exhibit a strong mylonitic fabric oriented parallel to foliation in their host rocks. All three of these samples yielded Late Jurassic magmatic crystallization ages (see Geochronology section below). The field and age relationships suggest that metamorphism and deformation of the Placerita Formation was ongoing during intrusion of the Late Jurassic granites.
Late Cretaceous Quartz Diorite (Map Unit Kqd)
All previously described map units are intruded by a medium-grained sphene biotite-hornblende quartz diorite pluton, informally named the Los Pinetos quartz diorite. This rock was mapped as part of the “Wilson Diorite” suite by Miller (1934), and later designated as “granodiorite” by Oakeshott (1958). It resembles multiple quartz diorite plutons in the central and eastern San Gabriel Mountains that have yielded U-Pb zircon crystallization ages between 82 and 71 Ma (Barth et al., 1995b; Nourse and Premo, 2016). The main quartz diorite body in the southern half of the map area exhibits a weak magmatic foliation defined by oriented biotite and hornblende. Pegmatite veins that crosscut the foliation are common (Fig. 17I). The Los Pinetos quartz diorite clearly post-dates the tectonic foliation preserved in common xenoliths of Placerita Formation, dioritic gneiss and Late Jurassic granite.
Two sills of moderately foliated quartz diorite intrude the Placerita metasedimentary rocks and Jurassic granite in Placerita Canyon upstream from the visitor center. The larger of these sills is injected by pegmatite and aplite veins, and in turn cut by abundant small faults and fractures associated with hematite, presumably related to the nearby Placerita fault of Battle (1939).
Two thin sections of Los Pinetos quartz diorite (samples JN 1808 and 1820, confirm that the foliation is generally magmatic rather than tectonic. These are the only sections examined in which quartz grains are not recrystallized. Medium to coarse anhedral quartz is slightly undulose, while pristine biotite and coarse hornblende grains display a preferred orientation that defines the magmatic flow foliation. A third thin section collected from the northern contact of the quartz diorite with Placerita Formation (sample JN 1821) shows modest subsolidus recrystallization of elongate quartz grains. We interpret this fabric to represent minor post-emplacement slip along the pluton margin.
Neogene Sandstone and Conglomerate
Crystalline rocks of the study area are bounded on the north by sandstone and granule-pebble conglomerate beds (map unit Tss-cg and Tcg), first mapped by Battle (1939) as the upper Pliocene or Pleistocene Saugus Formation, then reclassified as upper Pico Formation by Oakeshott (1958). We measured north-dipping bedding in this unit in cuts along Placerita Canyon Road (which is slightly mislocated on the topographic base map). The intervening Placerita fault juxtaposes these beds with vertically foliated biotite gneiss in Placerita Canyon (Fig. 16).
In the western part of the map area northeast-dipping Los Pinetos quartz diorite is overlain in angular unconformity by a gently (~20°) west-northwest dipping section of Saugus–upper Pico Formation. In several places we have distinguished a cobble conglomerate unit (Tcg) that occurs at the base of this section.
Quaternary Colluvium and Alluvial Terraces
Denudation of the area by wildfires revealed extensive landslide deposits, debris flows, and colluvium (mapped as Qcv) that occupy many of the ridgelines and obscure continuity of the bedrock units. The southeast branch of Los Pinetos Canyon also contains several subhorizontally bedded gravel terrace deposits (Qoa) in its upper reaches.
Limerock Canyon derives its name from exposures of white crystalline marble that occur in the lower reaches of that drainage. The area was highlighted by Oakeshott (1958) and Powell (1993) as an important place to study metasedimentary rocks similar to the Placerita Formation of Miller (1934). Both authors described metamorphic mineral assemblages in detail, and Powell proposed the name “Limerock Canyon Assemblage” in his correlations to a belt of metasedimentary strata dispersed along strands of the San Andreas fault (SAF) system. However, the only geological mapping in the area is Oakeshott’s rendition that shows a sharp contact along the trace of Limerock Canyon that separates a northwest-elongate body of Placerita Formation on the southwest from diorite gneiss with several limestone lenses on the northeast. Oakeshott mapped the metamorphic rocks as xenoliths within granodiorite of presumed Jurassic or Cretaceous age.
When Nourse first visited the area in fall 2016 to sample quartzite units, the canyon was choked with brush. Subsequently, the Creek fire of December 2017 denuded the entire canyon, and heavy winter rains of 2018 scoured away soil cover, opening access to some truly spectacular outcrops. Our geologic map (Fig. 18) shows essential elements of the metasedimentary series and associated intrusive rocks. Lithologies and metamorphic assemblages correlate very well with the rocks of Placerita and Los Pinetos Canyons. To avoid redundancy, we focus below on recently mapped field relationships, and describe new observations that constrain the timing of plutonism, metamorphism, and deformation.
General Characteristics and Map Relations
Limerock Canyon is carved approximately parallel to a north-northwest-striking, steeply dipping section of Placerita-type metasedimentary rocks. To characterize the stratigraphy, we mapped several quartzite and marble marker beds, 1–20 m thick, within a strongly foliated sequence composed of biotite gneiss, quartz-feldspar-biotite gneiss or schist, phyllite, and diorite gneiss (Fig. 18). Multiple dip reversals define map scale folds throughout the study area, indicating that much of the section is duplicated. One nicely exposed antiform is defined by a marble unit along the axis of the canyon below the third waterfall (Fig. 19A). The base of the section is defined by narrow bodies of granite augen gneiss basement, exposed in the core of two tight antiforms upstream from this waterfall. Stratigraphic continuity is broken by several east-northeast–trending faults with apparent left-lateral and/or normal displacements. Total structural thickness measured across strike is ~1200 m. Pervasive folding and concomitant ductile strain preclude determination of original thickness of the section as well as the bottom-to-top sedimentary succession.
Our mapping of Limerock Canyon is incomplete on the east wall, but the overall coverage contributes significant detail to previous works. Plate 1 of Oakeshott (1958) shows a 500-m-thick body of diorite gneiss with limestone lenses occupying the main axis of the canyon, and an irregular intrusive contact with granodiorite exposed part way up the ridge to the northeast. Our investigations indicate that most of Oakeshott’s diorite gneiss unit is biotite gneiss interlayered with quartzite and marble. Float derived from several side canyons is rich in biotite gneiss and quartzite clasts, indicating that Placerita metasediments are exposed far up the east canyon wall.
The crystalline rocks of Limerock Peak and Limerock Canyon are juxtaposed against Neogene sandstone and conglomerate of the Saugus Formation by the Lopez thrust fault (Oakeshott, 1958). This fault changes strike from east-southeast along the southern boundary of Limerock ridge to north-northeast where it crosses the lower part of Limerock Canyon. Several minor thrust faults with complementary and antithetic orientations relative to the Lopez fault cut the basement in lower Limerock Canyon. A major landslide shed south from Limerock Peak covers part of the Lopez fault.
New Field Observations
Paleoproterozoic Basement (Map Unit PCagn and PCgrgn)
The alluvium of Limerock Canyon contains large boulders of distinctive coarse-grained biotite granite augen gneiss. These strongly resemble Paleoproterozoic basement documented in the Mendenhall Peak area (Barth et al., 1995a), and in the eastern San Gabriel Mountains (Premo et al., 2007). They are also similar to components of the “Baldwin Gneiss” in the San Bernardino Mountains (Barth et al., 2000). We located two sources of this augen gneiss (mapped as PCagn on Fig. 18): (1) A cliff face in the wall of a side canyon draining the east side of Limerock Canyon has shed boulders in a restricted area of the lower canyon between the first and second waterfalls. That outcrop was inaccessible, but a large float block (sample JN 1926) from this canyon yielded prolific zircon. Later in our study we collected the along-strike continuation of this augen gneiss (sample JN 1931) from outcrop in the main eastern side canyon upstream from the third waterfall; (2) In upper Limerock Canyon, we mapped and sampled an exposure of augen gneiss (sample JN 1930; Fig. 19B) in the core of a tight anticline. Directly overlying beds of quartzite and biotite gneiss to the west form a tight syncline containing distinctive granule conglomerate and quartz pebble sedimentary breccia.
The augen gneiss outcrops constrain the deepest stratigraphic level of the metasedimentary series. Exposure 2 probably represents a paleo-unconformity, albeit strongly deformed and metamorphosed. This is an important relationship as it represents the first place in the San Gabriel Mountains where the Placerita Formation is observed to be in direct depositional contact with basement.
The augen gneiss of exposure 2 is interlayered on a decimeter scale with biotite gneiss that may be distinguished from similar gneisses of the Placerita Formation by its lack of cm-scale quartzite bands. Distinguishing the Proterozoic augen gneiss requires careful inspection because some of the Placerita biotite gneisses contain boudinaged layers of meta-arkose that display an augen-like appearance in high-strain zones.
A third exposure of Paleoproterozoic basement (unit PCgrgn) outcrops beneath a northeast dipping section of quartzite and graphitic schist on the north slopes of Limerock Peak. This rock is a strongly recrystallized, fine to medium-grained biotite-quartzofeldspathic gneiss (sample JN 1929). Dispersed augen of alkali feldspar 5–10 mm long suggest that the protolith was a slightly porphyritic biotite granite. 5-mm- to 2-cm-wide flattened and sheared pegmatite veins, oriented slightly oblique to foliation, indicate an original crosscutting vein network. The quartzite unit that directly overlies this gneiss becomes progressively coarser grained and more feldspathic approaching the contact. Northeast-dipping foliations are concordant across this contact that we interpret to be another occurrence of the basal unconformity (Fig. 18).
Limerock Canyon Assemblage
After Powell (1993), this term is applied to the metasedimentary rocks that we correlate with the Placerita Formation.
Graphite-Bearing Gneisses and Schists (Map Unitbgn). A distinguishing feature of the Limerock Canyon metasedimentary rocks is the presence of graphite as relatively concentrated (10%–25%) layers or as disseminated flakes within the biotite gneiss and biotite-quartzofeldspathic gneiss units. Oakeshott (1958) and Powell (1993) both emphasize this association and infer an origin from carbonaceous material entrained in the original sedimentary protoliths. Miller (1934) also mentions the presence of graphite in the type area of the Placerita Formation. We denote several graphite occurrences with gray squares on Figure 18.
All five thin sections cut from the biotite gneiss unit of Limerock Canyon preserve disseminated graphite that was not readily observable in hand specimen. The graphite forms dark blue-black opaque flakes and larger elongate shreds intergrown with brown biotite, defining the foliation in these gneisses. We also observed significant black graphite layers in a treacherous cliff face on the southeast side of Limerock Peak (Fig. 18). These exposures may represent the southeast continuation of graphite beds targeted by the McAnany and Rice mine (Oakeshott, 1958, p. 112) north of Limerock Peak. Our inspection of that graphite mine and a separate side canyon directly west revealed two distinctive horizons of graphitic schist, 5–10 m thick, within a section of quartzite and biotite quartzofeldspathic gneiss (Fig. 18).
Quartzite Marker Horizons (Map Unitsqandqcg). Our original intent at Limerock Canyon was to collect quartzite samples to compare detrital zircon signature with samples from Placerita Canyon and the eastern San Gabriel Mountains. New geological mapping delineates at least eight marker horizons containing 1–20-m-thick quartzite layers within biotite gneiss map unit (Fig. 19C). These rocks correspond to the “vitreous quartzite” unit of the Limerock Canyon assemblage described by Powell (1993). We collected five samples for detrital zircon geochronology to assess provenance variation within the Placerita Formation. Thin sections from these samples (FRC 2/18, JN 1605, JN 1632, and JN 1633, and JN1928; Fig. 18) preserve variable feldspar content (a few percent to 40%) and minor biotite, suggesting that some were derived from arkosic protoliths. Prolific zircon yields preclude any possibility of a chert origin. A granoblastic or hornfels texture reveals coarse quartz grains that are undulose and variably recrystallized.
We also mapped several occurrences of quartz-rich metaconglomerate (unit qcg) in proximity to the folded Proterozoic augen gneiss (PCagn) and medium-grained granite gneiss (PCgrgn) exposures in upper Limerock Canyon (Fig. 18). These beds are 20 cm to 1 m wide and vary in composition from quartz pebble sedimentary breccia to arkosic granule conglomerate (Fig. 19D). They are interlayered with phyllite and fine grained quartzofeldspathic biotite schists presumably derived from muddy and silty protoliths. We interpret the metaconglomerate outcrops to represent basal units of the Placerita Formation.
Mesozoic Intrusive Rocks
Early Jurassic Dioritic Gneiss (Map Unitdign). The lower part of Limerock Canyon displays important field relationships between the Placerita Formation and a “diorite gneiss” unit that Oakeshott (1958), described as follows: “Here dark gneissic dioritic rocks, high in hornblende or biotite, are in bands from a fraction of an inch to several inches in thickness interlayered with crystalline limestone, quartzite, and other metasedimentary rocks of the Placerita Formation. The more massive parts of the diorite can be seen distinctly cross-cutting Placerita metasedimentary rocks in some places and, in turn, light-colored granitic rocks sharply crosscut locally both the older units.”
We concur with Oakeshott’s description and add these details: Tabular bodies of fine- and medium-grained foliated diorite, a few cm to several meters wide, are interlayered with the biotite gneiss section containing quartzite and marble horizons. We sampled two phases: (1) JN 1923, medium to fine-grained hornblende quartz diorite; and (2) JN 1923B (Fig. 19E); fine- to very fine-grained biotite-hornblende schist. Both lithologies share a common tectonic foliation with the metasedimentary host, but a few less-deformed exposures also show sill-like intrusive relationships. A thin section of sample 1923 reveals that the quartz diorite protolith is strongly overprinted by a granoblastic fabric, with weak foliation expressed by elongate aggregates of hornblende and enhanced by aligned biotite that is completely replaced with chlorite.
A key field relationship is exposed directly south of the third waterfall, where a 4–8 cm layer of fine dioritic gneiss is isoclinally folded about a core of quartzite. This outcrop (Fig. 19F) demonstrates that the tight to isoclinal folding event observed throughout the Placerita sedimentary section is younger than the diorite.
Late in our investigation we mapped a thicker (~30 m) body of strongly lineated, medium-grained hornblende diorite gneiss along the axis of the canyon above the third waterfall. This rock forms a sill between Paleoproterozoic augen gneiss and Placerita Formation. Geochronology sample JN 1941 is in process.
Middle Jurassic(?) Pegmatitic Granite. A rock unit not present in the type area of the Placerita Formation but abundant in Limerock Canyon is a distinctively red weathered pegmatitic granite. Too small to map at the scale of Figure 18, this granite forms well-foliated several-cm- to ~1-m-thick sills within the layered sedimentary series. Books of coarse biotite and flattened K-feldspar define the foliation in this quartz-poor rock that may range compositionally from alkali granite to quartz syenite. In several places, the pegmatite occurs as intrusive sheets interlayered with and sharing foliation with the fine-grained phase of the diorite gneiss. Cryptic contact relations (Fig. 19G) indicate the pegmatite is the younger of the two intrusive rocks.
Late Jurassic Biotite Granite (Map UnitJbgr). Dikes of medium- to fine-grained biotite monzogranite sharply intrude the foliated rocks along northeast trends. The largest of these dikes is 20–30 m thick and dips ~40° north-northwest (Fig. 19H). This dike outcrops continuously from the main canyon at least 400 m northeast, suggesting that most rock exposures in the east canyon wall are part of the older layered series rather than younger plutonic rocks mapped by Oakeshott (1958). We observed several smaller granite veins with similar orientations cutting metasediments in the main trunk of Limerock Canyon. The biotite granite resembles the large Late Jurassic body (sample JNKV 1810; Fig. 16) that intrudes folded elements of the Placerita Formation in upper Los Pinetos Canyon. Our geochronology sample JN 1925 provides a complementary upper age constraint on the timing of metamorphism and deformation in Limerock Canyon.
Late Cretaceous Biotite Granodiorite (Map UnitKbgd). A 300 m wide pluton of unfoliated biotite granodiorite, largely covered by Quaternary colluvium, crops out directly northwest of Limerock Peak (Fig. 19I). Its western margin is a north-trending intrusive contact that cuts obliquely across the Placerita metasediments. The eastern boundary may be disturbed by a northeast-trending fault, but our mapping is incomplete in this area. A similar body of granodiorite that cuts across the metasedimentary section northeast of Limerock Canyon extends to the east edge of the map area.
Structure of Placerita Canyon and Limerock Canyon Study Areas
Our systematic mapping of primary and secondary structures and metamorphic fabrics shows that the Placerita–Los Pinetos Canyons and Limerock Canyon study areas share a common deformational and magmatic history. We describe below the structural features of the two areas from oldest to youngest, aided by stereonet plots that constrain the geometry of deformation and intrusion.
Pre-Late Jurassic Primary Structures
The earliest structures in both study areas are cryptic due to intense overprint of Late Jurassic metamorphism and tectonism. Compositions of the oldest layered rocks indicate existence of a marine sedimentary sequence composed of shale and siltstone with subordinate quartzose and arkosic sandstone and carbonates. Likely this represented a continental margin sequence, deposited on Paleoproterozoic granitic basement, presently exposed only in the cores of antiforms in Limerock Canyon. All primary sedimentary structures are obscured by metamorphic fabrics, and original bedding has been completely transposed into the regional foliation. Likewise, primary igneous contacts associated with the Early Jurassic diorite are transposed; these intrusions were likely emplaced as sills given that all contacts are parallel to compositional layering in the host metasediments.
Late Jurassic Metamorphism and Tectonism
The dominant structure in both study areas is a northwest-striking, steeply dipping foliation associated with upper amphibolite facies metamorphism. This metamorphic fabric is macroscopically displayed as a planar preferred orientation of biotite and/or hornblende generally parallel to boundaries between different layers of different composition. Flattened aggregates of feldspar, quartz, epidote, and garnet enhance this foliation. In thin section the fabric is granoblastic with individual minerals recrystallized to fine grain size. Orientation of the foliation differs significantly between the two study areas. Stereonet plots of the foliation poles show that the mean strike in Placerita and Los Pinetos Canyons is ~N70W, with predominant northeast dips between 60° and 90°, and subordinate steep southwest dips (Fig. 20A). Foliation data from Limerock Canyon display a more symmetric pattern, with a mean strike of ~N25W and dips of 70°–90° to the northeast and southwest (Fig. 20B). Also present in Limerock Canyon is a pervasive shallow lineation with northwest trends that coincides with fold hinges in the area.
Folding of the foliation and unit contacts is pervasive across the entire exposed sections in both study areas. Map scale fold hinges indicated on Figures 16 and 18 delineate consistent reversals in dip direction. These hinges trend ~N45–60W in Placerita and Los Pinetos Canyons, and ~N15–20W in Limerock Canyon. Statistical analyses of poles to foliation plotted on the stereonets yield cylindrical best-fit fold axes consistent with map scale folds; i.e., mean trend/plunge is N55W/40 for the Placerita-Los Pinetos data set (Fig. 20A) and S26E/16 for the Limerock data (Fig. 20B). Stereonet plots of individual fold hinges and axial planes (Fig. 20C) show complementary trend differences between the two study areas, with generally shallow hinges.
Other than strike, the general structural geometry is consistent between the two study areas. Upright folds with steep axial planes and interlimb angles between 15° and 45° were produced by a shortening direction of N30 E in Placerita–Los Pinetos Canyon and N65 E at Limerock Canyon. Altogether, the foliation, fold axis, and axial plane data indicate that metamorphism and folding occurred during one regionally pervasive tectonic event. The compressional structures at Limerock Canyon are rotated ~35° clockwise relative to those at Placerita–Los Pinetos Canyons. This difference might reflect local variation in the original convergence direction, or rotation of the Limerock Canyon structural block between the Late Cenozoic Lopez and Weir Canyon thrust faults. Such a large degree of rotation might require a significant component of oblique slip on these faults.
Field relationships in both study areas show that this strong tectonic fabric developed after intrusion of medium- to fine-grained diorite sills that share a strong foliation with their sedimentary host strata and are also tightly folded about northwest-trending hinges. A sample of diorite gneiss from Limerock Canyon (JN 1923B) yielded an Early Jurassic magmatic crystallization age of 180 ± 3.6 Ma. Along Santa Clara truck trail, granite gneiss sample JNKV 1807 was collected from the largest of several tabular intrusions into a fine-grained diorite gneiss body (Fig. 17H). Foliation in the granite is concordant to that in the diorite gneiss. This rock yielded a crystallization age of 151 ± 3 Ma. The tectonic fabric is sharply crosscut by unfoliated medium to fine-grained biotite granite in both areas. U-Pb zircon ages for these granites are 152 ± 3 Ma in Limerock Canyon (sample JN 1925) and 147 ± 3 Ma in upper Los Pinetos Canyon (sample JNKV 1810). Together, these relationships document an important Late Jurassic metamorphic and tectonic event.
Late Jurassic Synkinematic Intrusive Rocks
In Placerita Canyon, a strongly foliated and lineated tabular body of biotite granite (map unit Jgrgn) appears to be synkinematic with respect to deformation of the Placerita Formation. We observed granites of this type both as intrusive sheets into Placerita metasediments and dioritic gneiss, and as xenoliths within the Late Cretaceous Los Pinetos quartz diorite. While the Placerita Canyon granite displays a macroscopic mylonitic fabric in outcrop, thin sections reveal the texture to be statically annealed, presumably due to thermal effects of the Los Pinetos quartz diorite.
Sample JN 1722 was taken from a 100-m-wide sill of biotite granite that crosses the main trace of Placerita Canyon (Fig. 16). This granite gneiss displays a strongly lineated, steeply southwest-dipping mylonitic fabric that is locally discordant to foliation in the metasedimentary host rocks. It also contains a concordant xenolith of biotite gneiss whose foliation is rotated (N15E/56NW) relative to nearby biotite gneiss host rocks. Zircons from the Placerita Canyon granite gneiss yielded generally discordant 206Pb/238U ages between 111 and 171 Ma. Despite separate efforts to date this sample on the LA-ICP-MS and SHRIMP-RG, the results are difficult to interpret (see Tables 2 and 3; also Geochronology discussion below). Depending on which grains are excluded we calculated weighted mean 206Pb/238U ages between 145 and 128 Ma, with large uncertainties.
Late Cretaceous–Early Tertiary (Laramide) Fabric
As discussed later, Laramide age structures are locally manifested as a weakly developed subsolidous foliation in Late Cretaceous quartz diorite of the Placerita–Los Pinetos study area. Late Cretaceous granodiorites in Limerock Canyon do not appear to preserve Laramide fabric.
Late Cenozoic Faults
Several Late Cenozoic faults mapped during our study were not studied for kinematic information and are beyond the scope of this report. However, these faults are important in exposing the large blocks of Placerita Formation in both areas. The Placerita Canyon fault of Battle (1939) is likely a normal fault responsible for uplifting the type area of Placerita Formation in its footwall. Walrond (2004) interprets the Placerita fault as the southeast continuation of the late Miocene to pre–early Pliocene Palomas normal fault. Using well data, he shows that the Palomas fault has been overridden by reverse motion on the San Gabriel fault. The Placerita fault trends parallel to and runs just south of the San Gabriel fault.
The Limerock Canyon exposures represent a large block imbricated between the Lopez reverse fault and an un-named thrust fault. Also present in both areas are steeply dipping, northeast-striking left-lateral faults with minor displacements that are common in the San Gabriel Mountains as conjugates to strands of the San Gabriel and San Andreas fault systems.
U-Pb ZIRCON GEOCHRONOLOGY
Sampling Strategy and Analytical Methods
Between 2016 and 2019, we collected 25 geochronology samples to constrain the timing of deposition and metamorphism in the Placerita Formation. As mapping progressed, a consistent field sequence developed involving intrusions of presumed Mesozoic age that were either involved in the metamorphism or crosscut the related deformation fabric. Six samples have yet to be analyzed, but the current results summarized in Tables 2 and 3 provide essential information. Sample locations are shown on Figures 16 and 18.
Initial sampling focused on quartzites and meta-arkoses of the Placerita Formation, with the intent of obtaining detrital zircon populations to constrain provenance and maximum depositional age. We then collected multiple samples of foliated and unfoliated intrusive rocks to constrain crosscutting relationships and provide comparisons between the Placerita Canyon and Limerock Canyon study areas. Late in the study we discovered several bodies of augen gneiss in Limerock Canyon that constitute the basement upon which the sediments were originally deposited (although the nonconformity relations are now tightly folded and metamorphosed to upper amphibolite facies).
Zircons extracted from each sample were mounted in epoxy and imaged on our recently acquired Gatan Mini-CL instrument at Cal Poly Pomona. Four of the samples were analyzed on the SHRIMP-RG at Stanford University, using single-crystal analytical procedures for that laboratory (reference needed). The SHRIMP data were reduced using the SQUID 2.5 program (Ludwig, 2009). The remaining samples were analyzed at the LaserChron laboratory at California State University Northridge (CSUN) using the recently upgraded LA-ICP-MS instrument. Analytical procedures for the CSUN lab are described in in Part 1 above. Joshua Schwartz assisted with reduction and plotting this data.
Table 2 summarizes the principal results acquired at Stanford University, where several key intrusive rocks from Placerita Canyon and Santa Clara truck trail were analyzed. Table 3 gives results from the CSUN laboratory. Detrital zircon analyses are presented below with representative concordia diagrams (Fig. 21) and a series of probability density plots, filtered for discordance (Fig. 22 and 23). Below we describe the significant age constraints yielded by each major lithology.
Two samples of coarse-grained biotite granite augen gneiss from Limerock Canyon have been analyzed so far. A third sample of medium gained granite gneiss (sample JN 1929) is in process. Augen gneiss samples JN 1930 and JN 1931 both display similar discordia arrays, with upper intercept ages of 1741 ± 35 Ma and 1731 ± 35 Ma, respectively. Both preserve Late Jurassic lower intercept ages, likely related to the crosscutting granite intrusions known in this area. These samples represent the first known occurrences of basement in direct contact with depositionally overlying Placerita Formation.
Neoproterozoic Meta-Arkoses and Quartzites
Detrital zircon data from these rocks were frustrating to interpret because 56% of the 912 analyses were highly discordant, and many others were moderately discordant. Concordia diagrams of three typical samples (Fig. 21) show the majority of analyses scattered along crude discordia lines with late Mesozoic lower intercepts. Discordance may be attributed in part to lead loss related to Late Jurassic and/or Late Cretaceous thermal disturbances that are known in this area, but some of the data could also reflect discordance originally acquired in the source areas prior to erosion and deposition.
The probability density plots of Figure 22 show results after filtering the data for ≤6% reverse discordance and ≤30% normal discordance. Age maxima vary from sample to sample, but the most prominent peaks occur at ~0.7–0.8, 0.9–1.0, ~1.10, 1.35–1.45, 1.60–1.70, ~1.8, and ~2.0 Ga. Collectively the data set indicates several provenances that are well known in the nearby Mojave Province of southwestern Laurentia, although the 1.80 and 2.0 Ga populations are not characteristic. The Neoproterozoic grains are also anomalous in terms of known bedrock sources. A few scattered Late Archean grains, possibly derived from the Wyoming Province, were also analyzed.
To constrain maximum depositional age of the Placerita Formation, the youngest age population needs prudent interpretation. Most analyses younger than 1.0 Ga are highly discordant, but 28 grains from Placerita Canyon and 40 grains from Limerock Canyon survived the filtering criteria described above. A composite plot of 206Pb/238U ages (Fig. 23A) shows a significant age maximum between 900 and 1000 Ma, and a secondary maximum at ca. 780 Ma. To assess the possibility that these data represent thermally disturbed Mesoproterozoic grains, we calculated upper intercept ages for each analysis, using a lower intercept anchor that assumes lead loss at 80 Ma. The resulting plot (Fig. 23B) indeed shows a shift of many analyses to older ages (1.1–1.2 Ga; 1.4–1.5 Ga), however, a significant population remains at 900–1000 Ma, and a minor population at ca. 750 Ma. The Neoproterozoic grains are intriguing because it is difficult to tie them to western North American sources. Regardless of their provenance these results demonstrate a Neoproterozoic maximum depositional age.
Jurassic Diorite Gneiss
So far we have analyzed one sample (JN 1923B) from this rock unit; three additional samples are in process. The diorite is significant because it (1) represents the oldest generation of intrusion into Placerita Formation, and (2) intimately shares the deformational fabric (prograde metamorphic foliation and tight folding) observed in the host sediments. The resulting concordia age of 180 ± 4 Ma demonstrates that the Placerita Formation was metamorphosed and deformed after Early Jurassic time.
Late Jurassic Granites
We analyzed four samples of biotite granite to provide further constraints on the age of metamorphism. Two unfoliated samples (JNKV 1810 and JN 1825) that intrude sharply across metamorphic fabrics in the Placerita Formation yielded ages of 147 ± 3 Ma and 152 ± 3 Ma, respectively. Sample JN 1825 contains a Proterozoic inherited component consistent with the ages of wall rocks in Limerock Canyon. Several grains also record Late Cretaceous lead loss. A strongly foliated granite sill within diorite gneiss along Santa Clara truck trail (sample JN 1807 on Fig. 16) yielded a Concordia age of 151 ± 3 Ma. We interpret the fabric in this granite to record peak metamorphism in the Placerita Formation. Together with the early Jurassic age of the Limerock Canyon diorite gneiss, these results demonstrate that metamorphism and folding at Limerock Canyon occurred between 180 Ma and 152 Ma. The timing of metamorphism in the Placerita Canyon study area is more tightly constrained between 151 Ma and 147 Ma.
Another strongly foliated granite sample in Placerita Canyon (JN 1722) yielded a wide range of 206Pb/238U ages between 111 and 171 Ma (SHRIMP-RG). Mylonitic fabric preserved in this rock body shares a common orientation with foliation in the host Placerita Formation and diorite gneiss. The granite also contains foliated xenoliths of biotite gneiss and is intruded by Late Cretaceous quartz diorite. Poor statistics and pronounced discordance in the data from JN 1722 make a meaningful age difficult to resolve. A weighted mean 206Pb/238U age of 12 analyses with <5% discordance is 128 ± 7 Ma; while a weighted mean of all analyses yields 135 ± 5 Ma. These SHRIMP results did not improve much on widely scattered data produced during an earlier ICP-MS run, which yielded weighted mean 206Pb/238U ages of 136 ± 8 Ma from 13 near-concordant analyses and 145 ± 11 Ma from all 29 analyses. These Early Cretaceous ages (Tables 2 and 3) are unusual for the San Gabriel Mountains. We suspect that sample JN 1722 represents a Late Jurassic granite affected by Cretaceous lead loss, but there is little support for this thought in terms of elevated U/Th ratios in the younger grains. Nevertheless, it is interesting that similarly oriented mylonitic fabric is also developed in a nearby Late Cretaceous quartz diorite sill (Fig. 16). We believe that this part of the study area has experienced a Laramide age (Late Cretaceous–Early Tertiary) deformational overprint.
Late Cretaceous Intrusive Rocks
The Los Pinetos quartz diorite and Limerock Peak granodiorite record an important Late Cretaceous thermal culmination, yielding ages of 79 ± 1 Ma and 84 ± 3 Ma, respectively. These plutons sharply crosscut the metamorphic fabric in the Placerita Formation and are the most likely driver of profound lead loss in the detrital zircon assemblage. Lower greenschist facies foliation occasionally observed in the Los Pinetos quartz diorite is interpreted to represent a Laramide fabric that is only sporadically developed in the Placerita Canyon study area.
DISCUSSION AND REGIONAL IMPLICATIONS
In addition to describing detailed stratigraphic relationships for the first time, this article resolves long-standing questions about the age of the Placerita Formation, as well as the timing of metamorphism and associated intrusions (Miller, 1934; Oakeshott, 1958; Powell, 1993).
Significance of Paleoproterozoic Basement
Of special interest is the identification of 1741 ± 35 Ma and 1731 ± 35 Ma granite augen gneiss basement beneath a folded unconformity in Limerock Canyon. These are a rare basement age for the San Gabriel Mountains. Augen gneisses with ages between 1.69 and 1.67 Ga are well documented in the central and eastern San Gabriel Mountains (Premo et al., 2007) and across the San Gabriel fault within the Mendenhall gneiss (Barth et al., 1995a). We know of only two other occurrences of the older basement (Premo et al., 2007): a 1769 ± 10 Ma biotite granite gneiss from Cobal Canyon in the Potato Mountain block (Fig. 15), and a 1746 ± 8 Ma augen gneiss sampled from a float block in Cow Canyon landslide (derived from inaccessible exposures on western Ontario Ridge (Fig. 15). Both of these basement exposures are associated with thick packages of pelitic gneiss, quartzite, and marble that Powell (1993) correlated with the Placerita Formation. The association of ca 1.75 Ga basement with continental margin metasedimentary rocks of the Placerita Formation, together with absence of 1.67–1.69 Ga augen gneisses and pre–San Andreas fault reconstruction, suggests that Placerita Canyon, Limerock Canyon, Potato Mountain, and Ontario Ridge represent a distinctive Proterozoic crustal block that fringes the southern margin of the San Gabriel Mountains.
The Limerock Canyon granitic gneiss exposures are also important because they represent the only known place where basal conglomerates of the Placerita Formation are in direct contact with basement, although the unconformity has been tightly folded. The aforementioned 1.77–1.75 Ga gneisses of the eastern San Gabriel Mountains are actually isolated xenolith exposures within Cretaceous leucogranite and migmatite, respectively.
Depositional Age and Regional Correlation of the Placerita Formation
Our detrital zircon results, adjusted for 80 Ma lead loss, demonstrate a Neoproterozoic maximum depositional age for meta-arkoses and quartzites of the Placerita Formation. The youngest coherent age population is ca. 900–1000 Ma, although a few younger were also analyzed. At this point we are not sure if the protolith sediments were deposited during rifting of the Rodinia supercontinent at ca. 750 Ma (Whitmeyer and Karlstrom, 2007), or during later subsidence associated with the Cordilleran miogeocline (Stewart et al., 2001). We have considered possible correlations to the Pahrump Group and overlying strata of the Death Valley area (Mahon et al., 2014b). Many elements of the Placerita detrital zircon data set overlap with various Mesoproterozoic to Neoproterozoic stratigraphic units of Death Valley, however, noticeably missing from Mahon data set are the 1.8–2.0 Ga populations; also ages between 900 and 1000 Ma. Nevertheless, it is interesting that Mahon et al., (2014a) report Neoproterozoic ages for only 6 detrital zircon analyses out of 1098, but these near-concordant grains yielded a robust weighted mean 206Pb/238U age of 775 ± 18 Ma. In the Placerita Formation, the small population of youngest ages (seven grains between 760 and 840 Ma) overlap with the results from the Pahrump Group. Although most of our analyses were moderately discordant with large uncertainties, they may be statistically significant.
Closer to the San Gabriel Mountains, we are also exploring potential correlations to the Big Bear Group of the San Bernardino Mountains (Cameron, 1981; Barth et al., 2009). That quartzitic section unconformably overlies various components of the Baldwin Gneiss that preserve U-Pb zircon ages between 1760 and 1783 Ma (Barth et al., 2000). Detrital zircons from the basal Wildhorse Meadows quartzite unit display significant age populations of 1893–1785 Ma, 1496–1395, and 1257–1148 Ma, with two concordant Neoproterozoic analyses at 950 and 822 Ma (Barth et al., 2009). Two stratigraphically higher units preserve the three major populations noted above, with the addition of a significant 1786–1640 Ma population in the youngest horizon (Moonridge quartzite). The Placerita Formation compares reasonably well in that the 1.73–1.74 Ga basement is similar, and most of the detrital zircon populations from the Big Bear Group are represented. It seems especially noteworthy that the cryptic 0.9–1.0 and ca. 2.0 Ga populations we observe in parts of the Placerita Formation also show up as a subdued peaks in the Wildhorse Meadows data (Barth et al., 2009).
The possibility of exotic provenance for some of the Placerita detritus should also be considered, given that these rocks constitute a narrow belt of continental margin sediments situated at the western rifted edge of Laurentia. Workers interested in Rodinia reconstructions have long promoted the idea of continents attached to western North America during Mesoproterozoic time. Much debate persists (Moores, 1991; Karlstrom et al., 1999; Sears and Price, 2000) about which continent rifted away from the California margin at ca. 750 Ma. In a related discussion, Barth et al. (2009) present arguments for a western source for some of the detritus in the Big Bear Group. Similarly, although much of the Placerita Formation Proterozoic detrital signature can be linked to Mojave Province sources, persistent populations at 1.8–2.0 and ~0.9–1.0 Ga identified in our data have no obvious southwest North American provenance. Our search for a rifted “mystery continent” continues.
Jurassic Plutonism and Metamorphism
Other important contributions of our work include the identification of Jurassic plutonism in the study area, and new constraints on the age of prograde metamorphism and deformation. Consistent field relationships and geochronology of one key sample (JN 1923B) demonstrate that the earliest generation of diorite intrusions into the Placerita Formation occurred at 180 ± 4 Ma, prior to upper amphibolite facies metamorphism and associated folding. These diorites were previously recognized as pre-metamorphic by Miller (1934) and Oakeshott (1958), who speculated ages of early Mesozoic or late Paleozoic.
A Late Jurassic minimum age for metamorphism of the Placerita Formation is documented by granite intrusions in Limerock Canyon (JN 1925; 152 ± 3 Ma) and upper Los Pinetos Canyon (JNKV 1810; 147 ± 3 Ma). These granites sharply crosscut metamorphic fabric and folds in both areas. A well-foliated Late Jurassic granite sill (JNKV 1807) that shares a common fabric with its diorite gneiss host in the Placerita–Los Pinetos study area is 151 ± 3 Ma. These results imply that regional metamorphism occurred during Late Jurassic time. Our new data set also establishes the presence of Late Jurassic plutonism in the western San Gabriel Mountains.
Late Cretaceous Plutonism and the Laramide Orogeny
Two new ages for the Limerock Peak granodiorite (JN 1933; 84 ± 3 Ma) and the Los Pinetos quartz diorite (JNKV 1808; 79 ± 1 Ma) extend the known presence of Late Cretaceous magmatism to the western San Gabriel Mountains. These plutons are part of a batholith that includes plutons in the central and eastern San Gabriel Mountains that have yielded U-Pb zircon crystallization ages between 82 and 70 Ma (Barth et al. 1995b; Nourse and Premo, 2016). Similar plutons intruded into Proterozoic gneiss were originally mapped as “Wilson diorite” by Miller (1934) and “granodiorite” by Oakeshott (1958). Intense thermal activity associated with the Late Cretaceous batholith is known to have caused profound disturbance of detrital zircons in the Placerita Formation (this study) and quartzites of Ontario Ridge (Zylstra, 2017).
An additional occurrence of Late Cretaceous plutonic rock was recently documented by Barth et al. (2019) in the core of the Mill Canyon structural window (Carter, 1980) ~20 km northeast of Placerita Canyon. Here a 73 ± 2 Ma granite preserves a strong mylonitic fabric interpreted to be Laramide (Late Cretaceous-early Tertiary) age, based on Paleocene muscovite and biotite cooling ages. In the eastern San Gabriel Mountains, the Laramide orogeny is generally manifested by mylonitic rocks of the Vincent thrust system that include Late Cretaceous protoliths, including the 71 ± 1 Ma Mount San Antonio quartz diorite and a 70 ± 1 Ma leucogranite of Middle Fork Lytle Creek (Nourse and Premo, 2016). The intensity of fabric diminishes away from the Vincent thrust, but large areas of well-foliated Late Cretaceous quartz diorite persist in structural culminations as far west as the west fork San Gabriel Canyon.
In Los Pinetos Canyon, subsolidus fabrics of presumed Laramide age are locally developed along the northwest margin of the Late Cretaceous Los Pinetos quartz diorite. Stronger foliation is observed within a sill of similar quartz diorite along the main axis of Placerita Canyon, and mylonitic fabric pervades a nearby, poorly dated granite. These are the only places we have encountered fabrics of likely Laramide age. Farther southeast in Limerock Canyon, the Limerock Peak granodiorite is generally unaffected by Laramide overprint, although a weak magmatic foliation manifested by aligned biotite and is observed in thin section.
We conducted a geological mapping and U-Pb zircon geochronology study in Placerita, Los Pinetos, and Limerock Canyons to better constrain the depositional and tectonic history of the Placerita Formation. We found that the Placerita Formation is composed of Neoproterozoic continental margin marine metasedimentary strata, deposited upon ca. 1.78–1.70 Ga augen gneiss basement and intruded by Early Jurassic, Late Jurassic, and Late Cretaceous sills and plutons. Detrital zircons analyses from meta-arkoses and quartzites in the Placerita Formation are largely discordant but yield a Proterozoic age signature consistent with derivation from multiple Mojave Province bedrock sources. A minor Late Archean component probably records a distal provenance in the Wyoming Province. The youngest coherent age population is ca. 900–1000 Ma, demonstrating a Neoproterozoic maximum depositional age for the Placerita Formation. Sedimentary protoliths were subjected to upper amphibolite facies metamorphism concomitant with tight folding about northwesterly hinges during Late Jurassic time. This intense tectonic event is bracketed by foliated sills of diorite gneiss (180 ± 4 Ma) and granite gneiss (151 ± 3 Ma) that are crosscut by unfoliated biotite granites (152 ± 3 Ma and 147 ± 3 Ma, respectively). Profound Late Jurassic(?) and Late Cretaceous thermal overprint has caused lead loss in the majority of zircons from the Placerita Formation. This disturbance culminated with emplacement the Limerock Peak granodiorite and Los Pinetos quartz diorite at 84 ± 3 Ma and 79 ± 1 Ma, respectively. Effects of the Laramide orogeny are locally developed in the Los Pinetos intrusion.
ROAD LOG FOR PART 2
Drive the 14 Freeway north from the southbound Interstate 5. Take the Placerita Canyon Road exit, and turn right at the stop sign. Drive 1.5 miles to a right turn into Placerita Canyon County Park; park on the left near the visitor center.
Stop 2.1. Placerita Canyon—Day 1, ~2 p.m. (34.377960°, −118.467611°)
Walk east (upstream) in the main canyon the past visitor center to view a section of Placerita Formation first described by Miller (1934). Refer to the geologic map of Figure 16, as our traverse cuts obliquely across strike of most of the rock units. Please refer to Part 1 of this guide for detailed rock descriptions of this portion of our field trip.
The first outcrops in Placerita Canyon display biotite gneiss dipping northeast beneath a complexly folded interval of quartzite and calcsilicate gneiss, with minor biotite gneiss. This is a good place to demonstrate the sedimentary character of the Placerita Formation protoliths and upper amphibolite facies metamorphic assemblages. Farther upstream on the south side of the canyon, the section transitions to biotite gneiss with 20-cm- to 2-m-thick quartzite layers. Multiple dip reversals define several tight folds. Beyond here the metasediments generally dip northeast, with occasional dip reversals. An up-canyon transect reveals the following stratigraphic sequence of map units.
Biotite gneiss with quartzite layers intruded by a weakly foliated 40-m-thick sill of biotite hornblende quartz diorite (similar to dated Late Cretaceous sample JN KV 1808).
Fine- to medium-grained biotite granite sill with mylonitic fabric of variable intensity. Geochronology sample JN 1722 (135 ± 5 Ma?) was collected from a strongly foliated and lineated granite outcrop that contains screens of Placerita biotite gneiss.
A thick, coherent section of biotite gneiss, exemplified by the photograph in Figure 17A.
Quartzite and calcareous quartzite marker unit with local folds. This marker is crossed twice by Placerita Canyon and forms the southwest boundary of a thick Late Cretaceous quartz diorite sill. Geochronology sample JN 1720 (Fig. 17E) was collected from the second exposure.
Medium-grained, weakly foliated biotite-hornblende quartz diorite sill, similar to geochronology sample JNKV 1808 (79 ± 1 Ma). Several bends in Placerita Canyon transect extensive exposures of this 180-m-thick intrusion, cut by pegmatite and aplite granite dikes and hematite-striated faults likely associated with the nearby Placerita fault.
Intermittently exposed biotite gneiss with prominent white marble marker beds.
Sandstone and conglomerate beds of the Plio-Pleistocene Saugus Formation, juxtaposed against the Placerita Formation by the Placerita Canyon fault, poorly exposed in this area.
Return to vehicles and drive 1.6 miles east along Placerita Canyon Road to a locked gate on the right. Pass through gate (your leader has access key), and drive down into Placerita Canyon. Park near trailhead (Los Pinetos Canyon trail and Waterfall Trail). If a key isn’t available, the parking site is an ~300 m walk along a dirt road from Placerita Canyon Road.
Stop 2.2. Los Pinetos Canyon (34.376565°, −118.443978°)
Walk south up the main branch of Los Pinetos Canyon, staying in canyon bottom rather than climbing up Los Pinetos trail. Take the right fork where the canyon narrows.
Our traverse begins in biotite gneiss with several dip reversals, intruded by sills of unfoliated Late Jurassic biotite granite. About 120 m upstream from the first fork, north-dipping biotite gneiss is underlain by a thick section of calcsilicate gneiss with cm- to dm-wide quartzite and marble bands. This unit displays multiple folds at outcrop scale, and amphibolite facies mineral assemblages are well-displayed. One spectacular fold (Fig. 17D) may be seen by scrambling a short distance up the side canyon from the outside of the first left bend in the main canyon.
The main canyon narrows from this point on, where the calcsilicate section contains several irregular intrusions of foliated Jurassic diorite and amphibolite. Two north-northeast–striking faults are well exposed. Just before the next canyon junction is a sharp north-dipping contact with a biotite gneiss unit containing abundant thick quartzite layers. Take the right fork to view quartzite geochronology sample locality JN 1800. This outcrop is just downstream from a major waterfall that is impassable.
Return to vehicles and drive 3.2 miles west along Placerita Canyon Road, passing under the 14 Freeway. Turn right onto Sierra Highway, go 0.1 miles and take southbound 14 Freeway onramp. Drive 1.3 miles; exit at Newhall Avenue. Turn left, drive 0.2 miles to the parking area for hiking trails. Look to the right for a gated fire road (Santa Clara Truck Trail). Reset odometer here. Pass through the gate and follow the main dirt road. Angle right at the first branch, then left at the next branch. 2.1 miles up from the first gate is a second locked gate. Park on the right just up from the second gate.2
Stop 2.3. Neogene Nonconformity and Cretaceous Quartz Diorite (34.360978°, −118.473036°)
This roadcut displays a well-exposed west-northwest–dipping nonconformity between the Plio-Pleistocene Saugus Formation and Late Cretaceous biotite-hornblende quartz diorite. Two layers of cobble conglomerate occur just above the contact. Walk a couple hundred meters up the road to view a very fresh exposure of quartz diorite. This outcrop is approximately on strike with geochronology sample location JNKV 1808 (79 ± 1 Ma). A weak magmatic foliation is defined by aligned hornblende and biotite. Foliation is locally enhanced by elongate xenoliths of middle Jurassic(?) diorite gneiss (Fig. 17G).
Return to vehicles and continue up Santa Clara Road to pullout on outside of sharp left bend (Mile 2.9).
Stop 2.4. Foliated Late Jurassic Granite Dikes Intrude Middle Jurassic Diorite Gneiss (34.358719°, −118.467495°)
Here is the locality of geochronology sample JNKV 1807. This strongly foliated biotite granite (151 ± 3 Ma) is one of several sills intruded into middle Jurassic(?) diorite gneiss (Fig. 17H). Excellent exposures occur in a short walk down the road. The dated sample preserves a cryptic southwest-vergent mylonitic fabric that was found in thin section to be statically recrystallized.
Continue up Santa Clara Truck Trail, passing dated (79 ± 1Ma) quartz diorite sample locality JNKV 1808 at ~3.1 miles (34.360322°, −118.465763°). At ~3.6 miles (34.359360°, −118.456739°) is a prominent screen of Placerita Formation and Jurassic diorite gneiss within Late Cretaceous quartz diorite. Calcareous quartzite sample JNKV1809 yielded abundant zircon, but all analyses were extremely discordant due to Mesozoic thermal overprint. These stops are optional.
Continue up Santa Clara Truck Trail. At mile 3.9, a road that branches right (south) toward the microwave towers re-crosses the screen of Placerita Formation. Stay on Santa Clara Truck Trail (now officially Forest Route 3N17) and turn left at Wilson Canyon Saddle (mile 4.1). Primitive bathrooms are available at the saddle. At mile 4.4 is a gate on the left that marks the top of Los Pinetos Canyon trail. Stay on Route 3N17 and park at overlook on left at mile 5.1.
Stop 2.5. Undeformed Late Jurassic Granite Intrudes Deformed Placerita Formation (34.360603°, −118.437798°)
The intent of this stop is to show that undeformed Late Jurassic granite intrudes strongly metamorphosed and folded strata of the Placerita Formation. A walk down the road reveals two screens of biotite gneiss within weathered biotite granite. Also present is a thin dike of Late Cretaceous biotite-hornblende quartz diorite. The second screen contains several zones of folded biotite gneiss, exposed on a dirt road leading to a bee farm. A very fresh sample of this granite (JNKV 1810), collected along a slope southwest of the road just below the second screen, yielded a U-Pb zircon age of 147 ± 3 Ma.
Drive back down Santa Clara Road to the 14 Freeway, take southbound onramp. After ~3.7 miles, angle right onto eastbound 210 Freeway and return to Pasadena.
Day 2—Early a.m.
Drive the westbound 210 Freeway from Pasadena toward San Fernando. Take the Osborne Street exit. Turn right, then immediately turn left at the stop light onto Little Tujunga Canyon Road. Drive 3.4 miles to the mouth of Limerock Canyon. Park in the wide pullout on west side of road.
Stop 2.6. Limerock Canyon (34.312951°, −118.344404°)
The intent of this 90-minute hike is to give an overview of lithologies and crosscutting field relations in the type section of the Limerock Canyon assemblage (Powell, 1993). These rocks correlate with the Placerita Formation of Placerita Canyon and Los Pinetos Canyon. Please refer to the geologic map (Fig. 18) and detailed rock descriptions contained in the guidebook article.
Walk northwest up Limerock wash. Peruse the boulders that exemplify a wide variety of metamorphic and plutonic lithologies exposed in Limerock Canyon. Especially striking rock types are white crystalline marble for which this canyon is named, and coarse-grained biotite granite augen gneiss that constitutes 1.74 Ga basement upon which the Placerita Formation was deposited before metamorphism. Outcrops of Plio-Pleistocene Saugus formation outcrop on both sides of the lower canyon.
Where the canyon narrows, cross the late Cenozoic Lopez reverse fault with associated gouge and crush zones. One exposure in the wash juxtaposes biotite gneiss and coarse granite over Neogene sandstone. From this point the canyon cuts through a tightly folded section of biotite gneiss, quartzite, and marble, intruded by Middle and Late Jurassic sills and dikes. We will view the following outcrops and discuss new geochronology results.
Meta-arkose sample locality JN 1605, a typical layer of “quartzite” within biotite gneiss that forms the most abundant rock type of the Placerita Formation.
Diorite gneiss sample localities JN 1923 and JN 1923B. Two sills of biotite-hornblende diorite share a strong foliation with biotite gneiss and calcsilicate host rocks. Farther upstream near the main waterfall, similar diorite gneiss sills are tightly folded with Placerita Formation metasedimentary rocks. Here, the medium-grained diorite contains limited zircon, but the finer grained schistose phase (JN 1923B) yielded abundant apatite and zircon, from which we obtained a concordia age of 180 ± 4 Ma (“concordia age” is weighted mean 206Pb/238U age of overlapping concordant analyses that excludes other analyses). This sample provides a maximum age for metamorphism of the Placerita Formation protoliths.
White marble dips beneath vitreous gray quartzite. These rocks provide important marker units within the folded section.
Prominent northeast-dipping dike of unfoliated biotite granite crosscuts the Placerita metasediments. A fresh sample of this granite (JN 1925) yielded a concordia age of 152 ± 3 Ma. Together with the diorite gneiss, this sample constrains a Middle to Late Jurassic age for metamorphism.
Augen gneiss sample locality JN 1926. A fresh talus block derived from cliff exposures to the northwest yielded prolific zircons that have yet to be analyzed. The zircons have similar morphologies to those of dated samples JN 1930 and JN 1931 that yielded upper intercept ages of 1741 ± 35 Ma and 1731 ± 35 Ma, respectively.
Upstream is sequence of graphite-bearing biotite gneiss interlayered with thin quartzite layers. A prominent crosscutting vein of Late Jurassic(?) granite occurs at a small waterfall. We will turn around at the next waterfall, which requires a rope to scramble over.
A limited portion of the Placerita Formation is accessible from the second waterfall until one encounters a third waterfall that forms a major barrier. Rock units beyond this third waterfall (mapped on Fig. 18) may be accessed by walking southeast from the headwaters of Limerock Canyon.
Return to the vehicles and drive back to Pasadena. Stay on the eastbound 210 Freeway, passing cutoffs to the 605 and 57 Freeways. Take the Base Line Road exit. Turn left onto Base Line, then right at the next stoplight onto Padua Avenue. Please refer to road log for Part 3.
PART 3. LEGACY AND HAZARD OF LARGE LANDSLIDES IN SAN ANTONIO CANYON, SAN GABRIEL MOUNTAINS, CALIFORNIA (DAY 2, LATE A.M.)
Landslides of the San Gabriel Mountains
Southern California’s diverse rock types, topography, and climates produce landslides that range from mudflows to rock avalanches. Of these landslide types, long-runout rock avalanches and deep bedrock slumps have received the least study in southern California despite their abundance and area of landscape affected, probably because few have occurred in the brief historical window. Strong earthquake shaking (associated with M > 7.0 earthquakes) is the most likely trigger for these large landslides in seismically active regions with mountainous landscapes, such as the San Gabriel Mountains (Keefer, 1984, 2002; Keefer and Wilson, 1989; Morton et al., 1989). Importantly, mounting evidence suggests southern California has been in a major seismic energy release deficit (an “earthquake drought”) since ca. 1800 (Rockwell, 2016). A century of large earthquakes (and thus large landslides) may be on the horizon. Our last field trip segment focuses on large landslides in San Antonio Canyon of the eastern San Gabriel Mountains (Fig. 24).
In addition to the instant threat to communities and infrastructure in their paths (rock avalanches can travel many kilometers in minutes), many of these large landslide deposits have hazardous longer lived effects when they dam valleys (creating transient lakes upstream and a dam which may fail catastrophically) and erode to create significant influx of sediment to the valley downstream (which may cause the valley floor to rise tens to hundreds of meters). Dust clouds associated with coseismic landsliding in the 1994 Northridge Earthquake have even been shown to have caused a major outbreak of Valley fever (Jibson et al., 1998).
Landslides are traditionally mapped from topography by geomorphic features that are distinctly produced by landsliding (hummocky topography, headscarps, bulging toes, disturbances to drainage networks, variations in slope steepness, fracturing, differences in soil or vegetation, etc.). Landslide inventories are important, as past landslides inform us as to how and where future ones are likely to occur. The most recent inventory of landslide deposits in the San Gabriel Mountains was produced by the California Geological Survey (Bedrossian et al., 2012). That report is largely composed of mapping done by Morton and Miller (2006), compiled from 20-m-resolution topographic data, and remains incomplete in its spatial coverage. There have been considerable developments in the quality of topographic data available to interpret landslides in the last decade thanks to increasing availability of airborne lidar data (~1 m pixel resolution and the ability to image the ground surface beneath vegetation), which allows for unprecedented identification of landslide deposits (Fig. 25).
Three Quaternary dating techniques are primarily employed to date landslide deposits and related sedimentary deposits (underlying/overly deposits, impounded sediment upstream, aggradational terraces downstream) in the San Gabriel Mountains. Radiocarbon dating of charcoal or other entrained organic material is an ideal technique to utilize when the landslide deposit is suspected to be less than 50 k.y. due to the small sample size needed, cheap analysis, and low uncertainty ages but suitable material is rarely preserved. Single grain post-infrared infrared stimulated luminescence (post-IR IRSL; Rhoades, 2015) analysis of K-feldspars is well suited for dating clastic deposits in southern California where quartz grains often display low optically stimulated luminescence (OSL) sensitivity; in the right conditions this technique provides a burial age (~depositional age) for the grains in the sedimentary deposit. The most effective technique to directly date landslide events is using in situ-produced 10Be surface exposure dating (e.g., Ivy-Ochs and Kober, 2008) on large boulders on preserved surfaces of rock avalanche deposits. The deep-seated source, high deposit surface area to source area ratio, mixing boulder transport processes, and instantaneous deposition means that statistically it is very likely that the top surface of any given boulder began accumulating 10Be at the time of the landslide event (<10% of the boulders exhibit inheritance). This means that as few as three boulders can be sampled to yield a reliable age, in contrast to more typical exposure dating applications such as dating glacial moraines (in which every boulder likely has a different history and inheritance can be an issue).
Late Cenozoic Uplift of San Gabriel Mountains
The San Gabriel Mountains are part of the east-west–trending Transverse Ranges, located in a region of transpression along the San Andreas fault system in a semi-arid to sub-humid climate. The San Gabriel Mountains are primarily composed of fractured Paleozoic metamorphic and Mesozoic crystalline plutonic rocks (Morton and Miller, 2006) uplifted since 6 Ma by a range-bounding thrust fault system (Sierra Madre–Cucamonga fault system) (Matti and Morton, 1993; Blythe et al., 2002; Nourse, 2002). Many faults in the region are capable of M > 7 earthquakes, notably the Sierra Madre–Cucamonga fault (recurrence interval 600–700 yr), San Jacinto fault (recurrence interval 100–300 yr), and the San Andreas fault (recurrence interval ~150 yr) (UCERF-3).
Erosion rates increase with topographic relief (eastward increase across the San Gabriel Mountains; DiBiase et al., 2010). In the past 50 years virtually all erosion into the range-front basins is attributed to fire-flood-debris flow cycles (Lavé and Burbank, 2004). The drainage divide is highly asymmetric—85% of the San Gabriel Mountains drain southward toward the Los Angeles Basin, necessitating complex flood control infrastructure to protect the 12+ million population. An argument could be made that in terms of potential downstream effects the San Gabriel Mountains are one of the most important mountain ranges in the country. The Los Angeles Basin itself is a considerable sediment sink with a maximum depth of ~8 km (Wright, 1991; Magistrale et al., 1996).
Crystal Lake and North Fork San Gabriel River
A notable paradigm shifting case study by Scherler et al. (2016) examines extensive fill terraces in the North Fork of the San Gabriel River valley in the central portion of the San Gabriel Mountains (Fig. 26), which had previously been linked to changes in regional climate. This study provides the first cosmogenic exposure ages for some of the largest landslides in the San Gabriel Mountains, including the Crystal Lake Landslide, previously inferred to be 1–2 Ma but clearly shown to be ca. 4 ka. The Alpine Canyon Landslide occurred as recently as 600 years ago. Following one of these major landslide events Scherler et al. (2016) find evidence for incision rates (vertical downcutting) into the landslide deposits on the order of 35 mm/yr (high!), which in the subsequent 700 years filled the valley downstream with up to 150 m of sediment. Such high rates of sediment production and accumulation have important implications for the Los Angeles region following a major earthquake, with the potential to impact reservoirs, flood control, utility infrastructure, and homes. Together these indicate large landslides in the San Gabriel Mountains are an underappreciated and timely hazard worthy of further research. They are probably earthquake triggered.
Following on from the Scherler et al. study and with recently acquired lidar data in hand, University of California at Riverside M.Sc. student Chris Gentile and author Nicolas Barth began to develop a case study in San Antonio Canyon, which contains some of the highest relief in the San Gabriel Mountains and numerous valley-damming landslide deposits. The San Antonio case study is used as a jumping-off point for the ultimate goal of systematically mapping the entire San Gabriels for landslides, spatially analyzing trends, and unraveling temporal trends.
San Antonio Canyon
San Antonio Canyon drains the south flank of the highest peak in the San Gabriel Mountains, Mt. San Antonio (also known as Mt. Baldy, 3038 m), and is one of the few south draining catchments that abuts the main divide of the range (also San Gabriel River, Big Tujunga River). The total catchment area is 73.1 km2 to the flood control dam at the San Gabriel Mountains range front. San Antonio Canyon contains some of the highest relief and steepest slopes in the San Gabriel Mountains; it also contains a notable concentration of large complex slump and rock avalanche type landslides that have been of interest to many previous workers (Ehlig, 1958; Morton et al., 1989; Nourse et al., 1998; Morton and Miller, 2006).
Bedrossian et al. (2012) provide the most recent compilation of landslides in San Antonio Canyon, itself based largely on mapping by Morton and Miller (2006). Bedrossian et al. (2012) mapped 9.4 km2 of landslide deposits (n = 87). With detailed field mapping and newly acquired lidar data, Barth and Gentile mapped 18.5 km2 of landslide deposits (n = 264). There are three times as many landslides in San Antonio Canyon as previously mapped. The San Antonio Canyon catchment is 25% landslide deposits by area, twice the amount previously mapped (Fig. 27).
PART 3 ROAD LOG AND STOP DESCRIPTIONS
This road log picks up where the log for Part 2 terminated at the junction of Base Line Road and Padua Avenue in Claremont.
A side note: Google has complete 3D “Street View” coverage of the entire Part 3 field trip route (other than side hikes). Go to google.com/maps, enter the locations, and drag-and-drop the orange figurine in the bottom right corner onto the map to revisit every roadcut and vantage featured in Part 3.
Drive up Padua Ave. Notice abundant use of boulder masonry and landscape work locally derived from the San Antonio Canyon alluvial fan (debris flows and floods).
Turn right onto Mt. Baldy Road (Set odometer to 0).
While driving up the grade, notice the San Antonio Canyon alluvial fan surface down to the right, engineered debris basins on the left to protect the neighborhood of $1.5–3 M houses, and the roadcuts through red weathered terrace deposits on the right. The latter old terrace deposits are ~3 m above the base of San Antonio Dam. Similarly weathered terrace deposits on the San Gabriel Mountains range front in Pasadena (Jet Propulsion Lab campus) have a ca. 200 ka cosmogenic burial age (Burgette et al., 2018).
Stop 3.1. Mile 1.3: San Antonio Dam
San Antonio Dam was built by the Army Corps of Engineers in 1956 to prevent a repeat of the high waters that inundated Upland, Claremont, and La Verne in March 1938. It played a crucial role in reducing damage during the 1969 flood. In addition to flood control, the dam also serves the purpose of groundwater recharge. Its porous and permeable substrate supplies renewable well water to the City of Upland. San Antonio Reservoir has never exceeded the height of the spillway (2238 ft), but its water elevation reached 2193 ft and 2226 ft during January of 1969 and February of 1980, respectively. The earth-fill design of San Antonio Dam is prudent given its location near the intersection of three active faults that conceivably might rupture a concrete structure. As we drive up Mt. Baldy Road, notice the roadcuts through San Antonio Canyon alluvial fill at a higher elevation than the dam. These deposits have post-infrared infrared stimulated luminescence (post-IR IRSL) burial ages of 28.1 ± 2.0 ka and 21.7 ± 3.1 ka (Barth unpublished data); look across the canyon and consider the magnitude of downcutting and volume of sediment that must have been removed since this time.
Stop 3.2 (Mile 3.8). Evey Canyon Fault Outcrop (Fault Rock Textures)
This exposure of the northeast-striking Evey Canyon fault (itself a splay of the left-lateral San Antonio Canyon fault) provides one of our best opportunities to examine a fault zone in San Antonio Canyon (Fig. 28). Here it is seen as a northwest-dipping plane separating green and reddish-purple sheared rock. This fault zone has accommodated at least 3 km of left-lateral slip since the Late Miocene (Nourse, 2009).
Observe the style of fractures in the surrounding quartz diorite and in the zone of sheared rock, as these will be contrasted with rock avalanche textures at the next stop (Stop 3.3). Generally, most fault slip here is localized to a narrow zone dominated by discrete shears and relatively low-porosity cataclasite and clay alteration.
Stop 3.3. Mile 3.9: North of Spruce Canyon Rock Avalanche Outcrop (Rock Avalanche Textures)
This roadcut exposes the basal contact of a rock avalanche deposit with underlying sediments and is a good place to observe and discuss rock avalanche textures (Fig. 29). These textures are often contradictory: appearing both coherent and incoherent, exhibiting both brittle and ductile textures.
Crackle breccia: a transported block or zone of pervasively shattered rock in which the internal fracture-bounded fragments show little to no separation and rotation relative to one another and in which relict lithologic fabric is still discernible.
Jigsaw breccia: internal fragments are separated from one another by thin bands of matrix material but are not significantly rotated with respect to one another so that relict bands and fabric are traceable and continuous across the breccia, and the pieces can be “put together” as in a jigsaw puzzle. Color banding due to lithologic variation often a key trait.
Yarnold and Lombard (1989) examined about a dozen large rock avalanches in the southwestern United States to develop a facies model. Generally speaking, the deposits are inversely graded and show upward decreases in matrix percentage. From bottom to top they describe a characteristic stratigraphy:
Substrate: Upper surface of the substrate may be scoured to depths of 5–10 m. Bedding beneath the basal contact may be folded or contorted.
Basal contact: Can be planar and striated (like a thrust fault plane) or have an ultracataclastic layer and/or have a mixed zone (see below). Underlying pebbles can be striated or sheared off.
Mixed zone: Usually less than 5 m thick (may not be present). A discontinuous silty to sandy conglomerate predominated by substrate clasts but includes rock avalanche–derived material. The mixed zone can be homogeneous or contain internal bands or lenses, each with its own distinct texture, color, and lithology. Clastic dikes may extend upward from the mixed zone up to 3 m.
Breccia sheet: Consists of a disrupted zone near its base and an overlying matrix-poor zone. Locally contains well-developed slip surfaces.
Disrupted zone: Up to 10 m thick. Highly comminuted near its base. Matrix is dominantly derived from adjacent clasts. The disrupted zone grades upward into matrix-poor breccia.
Matrix-poor zone: Comprises most of the deposit’s thickness. Jigsaw and crackle breccia textures predominate. Blocks can be from a few meters to many tens of meters (“megabreccia”). Most blocks are internally brecciated, although some are surprisingly intact and strongly resemble in situ outcrops. Matrix is derived almost solely from adjacent clasts.
Boulder cap: Coarse boulders (>1–30 m across) comprise the surface of rock avalanche deposits in a loose, chaotic framework (it may be possible to climb down several meters in the void space between boulders). Despite the coarse grain size, hummocks and longitudinal ridges are often preserved in the geomorphology.
Outcrops such as this one are often misinterpreted as either faulted bedrock or angular clastic sediment. When freshly exposed, there is a tendency for the field geologist to misattribute the deformation as fault-generated, probably exacerbated by the fact that landslide-produced rock textures are rarely taught in structural geology curriculum. While catastrophic landslide and seismic fault movement can occur at similar slip rates (e.g., >1 m/s), a key difference is that landslide deformation occurs at the surface with low confining pressure, which leads to more diffuse deformation, greater mixing, and higher porosity. When outcrop quality is mediocre and partially covered by surface wash, the more resistant clasts stand out and the deposit can overall appear as an angular clastic sediment such as colluvium. Examining large clasts for crackle breccia textures and searching the matrix between for jigsaw breccia textures is a good test, since neither is suitable to form in a colluvial setting. Just as any initial hand sample classification should start with the question “Is it igneous, sedimentary, or metamorphic?”; in mountainous regions, examination of any outcrop should start with the question “Is it in situ bedrock, landslide deposit, or sediment?” When approached objectively, the answer to the latter question can often be ambiguous.
Stop 3.4. Mile 5.5: View of Spring Hill Rock Avalanche
West of the road is yet another outcrop of rock avalanche overlying sediment. Our attention here will mostly be focused on the Spring Hill Rock Avalanche deposit across the valley to the east (Fig. 31). Spring Hill is one of the older rock avalanche deposits preserved in the valley. Boulders are rare on the low gradient surface of the deposit, likely indicating poor preservation of the landslide’s boulder cap. The boulders present are subrounded to well-rounded and have a significant weathering rind with clear evidence for spalling from fires. Soil development on the surface is considerable; vegetation is grasses and chaparral. The geomorphic surface slopes gently toward the west as an apron but includes remnant hillock morphology; it is deeply incised by three major drainages including Cascade Canyon. The deposit is consolidated enough that it forms 65 m vertical cliffs with red weathering. From our vantage at Mt. Baldy Road, the contact between the cliff-forming rock avalanche deposit and the underlying greenish-gray Proterozoic layered gneiss can be clearly seen.
A minimum age derived from 10Be exposure of boulders is pending but an estimate of 45 ka is suggested based on similar preservation to a San Andreas fault offset rock avalanche in the Cajon Pass area to the east (S. McGill, 2020, personal commun.). The Spring Hill deposit was mapped as “Quaternary Older Gravels” by Dibblee and Minch (2002). It is the author’s experience that many Dibblee “Qog”s are in fact similarly old landslide deposits.
From this stop on the road the Spring Hill Rock Avalanche deposit appears to be a thin veneer deposited on a bedrock plateau extending over 800 m into San Antonio Canyon. However, our mapping indicates the deposit buries considerable paleo-topography and is 250 m thick in places. Together, the Spring Hill Rock Avalanche and the Cow Canyon Saddle Rock Avalanche to the north completely bury a 5-km-long stretch of San Antonio Canyon. In this area, San Antonio Canyon formerly followed the San Antonio Canyon fault before being buried by the landslides. Rather than downcutting through the landslide deposits to regain its former valley bottom, San Antonio Creek cut a new valley at a position ~700 m to the west. This is a major example of epigenetic gorge formation (when channels that have been laterally displaced during episodes of river blockage or aggradation incise down into bedrock spurs or side-walls of the former valley rather than excavating unconsolidated fills and reinhabiting the buried paleovalley). At Stop 3.6 we will see an active (smaller) example of epigenetic gorge formation due to the Hogback Rock Avalanche.
Stop 3.5. Mile 5.8: Hogback Rock Avalanche (Top)
A short walk to the southeast leads onto the well-preserved bouldery surface of the deposit.
The Hogback Rock Avalanche is one of the more significant young canyon-altering events in San Antonio Canyon. Notice the characteristic chaotic texture of the deposit’s “boulder cap,” the lack of soil development, and sharp edges of the boulders (indicating little erosion since deposition). The source area can be clearly seen to the northwest. This deposit is an ideal candidate for 10Be cosmogenic exposure dating (dates pending). At this elevation the vegetation should be manzanita scrub, but instead we have Bigcone Douglas Fir (up to 500 years old here) and Canyon Live Oak on the deposit, the latter two of which prefer steep, well-drained slopes and have deep root systems able to find water (Morton et al., 1989). The rocky texture and distinct vegetation are key clues to identifying many other landslide deposits in the canyon.
Notice the Hogback Rock Avalanche and the more complex slump-style landslides to the west (Fig. 32).
The modern road crossing the top of the Hogback Rock Avalanche was built in 1950. The “old” road (our next stop) cut into the margin of the Hogback Rock Avalanche was finally abandoned in 1965, after the first major rain in 27 years caused severe damage to it (Morton et al., 1989).
As we drive past the roadcuts through the Hogback Rock Avalanche deposit, notice that sections appear to be incredibly coherent. Dibblee and Minch (2002) mistakenly mapped these roadcuts as bedrock but they are in fact megaclasts within the rock avalanche deposit.
Stop 3.6. Mile 6.1: Hogback Rock Avalanche (Bottom)
A short walk on the old road to the southeast reveals the internal structure of the Hogback Rock Avalanche and the youthful bedrock slot canyon that is actively being cut.
Walk on the old Mt. Baldy Road cut into the northern margin of the Hogback Rock Avalanche (Fig. 33). Notice the large boulder clasts among the oaks on the right. On the left, notice the wider floodplain and multiple channels of San Antonio Creek. The Hogback Rock Avalanche spanned the full width of the canyon, damming San Antonio Creek long enough to create a lake upstream, which deposited fine-grained bedded sediments visible on the east side of the canyon. As the creek cut through and around the rock avalanche deposit, it cut a spectacular little slot canyon (an epigenetic gorge) into adjacent bedrock on the east side of the canyon to try to regain its former graded profile. There are sizable bedrock waterfalls below us here (the creek is still very much adjusting to the Hogback–Spring Hill Rock Avalanche). A strath terrace with overlying sediment is visible at the elevation of the road (an attempt at a post-IR IRSL age failed, possibly because of very rapid deposition and no bleaching). The eroding scree slopes above the old road bed contain some nice exposures of the internal structure of the Hogback–Spring Hill Rock Avalanche that can be accessed with care.
Stop 3.7. Mile 8.1: Cow Canyon Saddle
The parking lot at Cow Canyon Saddle is a great location to discuss the Cow Canyon Saddle Rock Avalanche, implications for major drainage reorganizations, and some recently dated “valley fill deposits” in Cow Canyon (a tributary of the San Gabriel River). The parking lot is built on the distal portion of the Cow Canyon Saddle Rock Avalanche (Fig. 34).
The 2.6 km2 Cow Canyon Saddle Rock Avalanche is the second largest rock avalanche deposit in the San Gabriel Mountains by area (after Crystal Lake). The rock avalanche traveled laterally up to 5 km away from a source near Cucamonga Peak on the east side of San Antonio Canyon. The Cow Canyon Saddle Rock Avalanche either (1) buried an existing low-relief saddle between Cow Canyon and San Antonio Canyon (<150 m), or (2) drove a major drainage reorganization in which 40 km2 of the upper San Antonio Canyon catchment was diverted from Cow Canyon to lower San Antonio Canyon. Regardless of whether the Cow Canyon Saddle Rock Avalanche is implicated, there is considerable geomorphic evidence that such a drainage reorganization may have occurred (regional drainage patterns, anomalously low channel gradient in Cow Canyon more consistent with a large upstream drainage; anomalously narrow and sinuous channel; Ehlig, 1958; Morton et al. 1989) (Fig. 35). Two 10Be ages for exposed boulders on the distal portion of the Cow Canyon Saddle Rock Avalanche (4097 ± 478 yr, 3981 ± 503 yr) have been collected and two more ages are pending.
Del Vecchio et al. (2018) obtained a 33–40 ka post-IR IRSL depositional age for southwest sloping deposits inferred to be emplaced by dry colluvial and debris flow processes. They note these surfaces have “soils exhibiting a greater degree of weathering than nearby soils formed on bedrock” and suggest a 1–4 ka soil age from regional chronosequences. They suggest that an apparent 10× difference between soil and depositional ages for deposits in Cow Canyon may be strong evidence for frequent soil stripping in response to wildfire or strong precipitation events.
The toe of the Cow Canyon Saddle Rock Avalanche is at a lower elevation than the southwest-sloping fill deposit beneath Lookout Mountain, suggesting the bulk of the fill deposit degradation occurred prior to 4 ka. So large landslides sourced from the north side of Cow Canyon occurred, followed by 33–40 ka colluvial aprons that buried Cow Canyon, that were later downcut by ~150 m prior to 4 ka, followed by the Cow Canyon Saddle Rock Avalanche at 4 ka. The valley-spanning fill deposits would be much easier to create if Cow Canyon was already beheaded of upper San Antonio Canyon. The valley-spanning deposits, however, would be much easier to incise down to current level with greater upvalley stream power. The landslide from Lookout Mountain could have cut Cow Canyon from its San Antonio Canyon head. The paleo-valley beneath Spring Hill Rock Avalanche and the newly formed valley could factor into the history. There should be enough evidence preserved to come up with a unique sequence of events; the chronology is still being pursued. Regardless, this is a lot of landscape change over a relatively short interval of time.
Return back down to Mt. Baldy Road and turn left (north) onto the road to continue the field trip. Drive through Mt. Baldy Village (population 500), which sits in a precarious position beneath steep slopes prone to rockfall (scree) and debris flows, forest fires, and a flood-prone creek bed. January 1969 saw 50 inches of rain in just 8 days in January, followed by more high-intensity rain in February, which resulted in over 100 houses lost in Mt. Baldy Village. The Buckhorn Restaurant and Lodge and the trout ponds are located at a site where the 1938 flood destroyed a flourishing Camp Baldy Resort, operated by a member of the family that established Yosemite’s Camp Curry (Morton et al., 1989).
Stop 3.8. Mile 9.6: Cow Canyon Saddle Rock Avalanche Exposure
Be careful on the road at this blind corner.
At this location, the base of the Cow Canyon Rock Avalanche deposit is below the elevation of the road, possibly indicating San Antonio Creek has not yet cut back to its former base level here (Fig. 36). At a distance, this outcrop could be easily misinterpreted as a fault zone, but looking closely, crackle breccia and jigsaw breccia textures are abundant where exposure is fresh. Breccia shear bands and lenses with distinct texture, color, and lithology are also distinct. The deposit is ~2 km away from its source at this location. The top portion of the outcrop appears to be alluvial material—it is in fact a part of the rock avalanche deposit too (notice more matrix-poor sheared rock above).
Continuing upcanyon, notice an abundance of scree slopes and debris flow-laid boulders. Continue left on Mt. Baldy Road past the junction with Icehouse Canyon.
Stop 3.9. Mile 12.8: Manker Flat Rock Avalanche
Manker Flat Rock Avalanche (“Mt. San Antonio Landslide” of Morton et al. 1989) is sourced on the southern flanks of Mt. San Antonio (also known as Mt. Baldy), the highest peak in the San Gabriel Mountains at 3068 m (Fig. 37). The Manker Flat Rock Avalanche deposit consists of Late Triassic Mt. Lowe quartz monzodiorite and Cretaceous quartz diorite (Nourse et al., 2010), distinctive from the bedrock lithologies juxtaposed along most of the landslide’s flow path. The distal portion of the landslide (another Dibblee and Minch 2002 “Qog”) is well-preserved with a stout snout and flow textures visible in lidar. The deposit is over 100 m thick in places. Its more proximal deposit area has been dissected by creek downcutting and remains as a series of benches on the east side of the canyon. 10Be dates are pending for distal and proximal portions of the Manker Flat Rock Avalanche, but it appears to be younger than the Cow Canyon Saddle Rock Avalanche based on geomorphic preservation.
The Manker Flat Rock Avalanche had a runout of over 4 km and traveled at great speed. As it came down upper San Antonio Canyon and crossed Manker Canyon, it “ran up” over 100 m in height up the other side of the valley. A portion of the rock avalanche then stalled on the hillside as the main mass continued downcanyon. Caution is necessary to avoid misinterpreting features as being fault scarps or related to multiple landslide events. In this instance the geology assists the interpretation since the igneous clasts of the rock avalanche deposit are notably different from the layered gneisses they are deposited on.
There may be a considerable volume of sediment impounded behind the rock avalanche deposit in Manker Canyon. Headward erosion is currently cutting through the rock avalanche deposit; erosion rates may increase once the knickpoint reaches the impounded sediment fill.
Several fresh and well-integrated arcuate scarps along the rim of Baldy Bowl and at the summit of Mt. San Antonio suggest a similar event could recur. Interestingly, most of these scarps indicate failure of catchment ridgelines; many large landslides in the San Gabriel Mountains include peaks or ridgelines and when they fail they increase the area of the catchment they are deposited into.
PART 3 SUMMARY
The largest landslides in the San Gabriel Mountains are late Quaternary–Holocene in age (not early Quaternary as previously thought). They should be considered a significant (if as yet unquantified) hazard associated with future earthquakes and are an important driver of Holocene landscape evolution. Large catastrophic landslides present not only the immediate hazard of the landslide itself, but also can have effects lasting hundreds of years (drainage impoundment, lakes, dam-break floods, heightened erosion/aggradation, etc.) or longer (drainage reorganization, epigenetic gorge formation). New mapping in San Antonio Canyon reveals three times as many landslides as previously mapped and that landslides cover twice the surface area previously mapped.
The best approach to mapping landslides in this landscape is to take a two-prong approach: (1) GIS mapping with lidar data to provide the most detailed view of geomorphology, and (2) field mapping with a structural geology focus to characterize rock textures and map contacts. Early Holocene landslides may already have their geomorphic expression removed. Once sufficient soil has developed and vegetation has colonized a rock avalanche surface, wildfires are highly effective at degrading the geomorphic surface. Older surfaces may look similar to gravel fill deposits. Accurately mapping the landslide extent can be very important for working out fault history, paleotopography, and drainage reorganizations.
Part 1. Swanson would like to thank Chris Wills and Jeremy Lancaster (retired and current California Geological Survey Mapping Program Managers) for making funding available to conduct the geochronology and geochemistry work discussed herein and for supporting continued study of this area. He would also like to thank his STATEMAP coauthor Brian Olson, who focused on other important aspects of the referenced quadrangle mapping and carried the heavy load on the digital aspects of map preparation, and Carlos Gutierrez (California Geological Survey) who was instrumental in preparing the geologic quadrangle maps for publication and who prepared the map figures for this section of the guidebook. Sadie Knepper and Anita Carney also assisted with figure preparation.
Part 2. Ranger Russ Kimura of the Placerita Canyon Visitor Center provided permits to access closed areas of the park. Nourse is grateful to Mike Dykstra, Mark Thompson, Frank Wille, Clark Murphy, and Larry James Martin for assistance with geological mapping in Limerock Canyon. Zircons were separated, mounted, and imaged using facilities housed in the Cal Poly Pomona Geology Department. Vanessa Pena helped with the mineral separation. The Teacher-Scholar program at Cal Poly Pomona provided funds to analyze the zircons at Stanford University and California State University Northridge. Efforts of Joshua Schwartz in reducing and plotting the LaserChron data are much appreciated.
Part 3. Barth would like to acknowledge the M.Sc. work of University of California at Riverside student Chris Gentile (which includes contributions to mapping and Quaternary dating referenced herein), helpful discussions with Doug Morton, Pete Sadler, Rich Minnich, Devin McPhillips, Janis Hernandez, Brian Swanson, and Katherine Kendrick, and lab assistance from John Southon, Seulgi Moon, Nathan Brown, Dick Heermance, and PRIME lab. Previous field guides to the San Antonio region (notably Morton et al., 1989; Nourse, 2009; Nourse et al., 2010) formed a useful starting place for this field guide’s Part 3 itinerary.