New mapping in the Lake Kaweah pendant of the southwestern Sierra Nevada batholith reveals a previously unrecognized nonmarine sequence of metamorphosed sedimentary and volcanic strata, defined herein as the Goldstein Peak Formation. The nonmarine origin distinguishes the Goldstein Peak Formation from all other Sierra Nevada metasedimentary pendants and from virtually all other coeval deposits associated with the Sierra Nevada arc. Basic structural relations, supplemented by new U-Pb zircon ages, suggest an Early Cretaceous depositional age, a time that is poorly represented within the stratigraphic record of California. This unusual age makes the Goldstein Peak Formation the youngest sedimentary deposit preserved within the metamorphic framework of the exhumed batholith, one that was deposited concurrently with some of the earliest deposits in the Great Valley forearc basin and just preceding the mid-Cretaceous Sierra Nevada arc surge. Preserved sedimentary and volcanic structures, along with whole-rock geochemistry, are consistent with deposition of Goldstein Peak conglomerates and sandstones within fluvial and alluvial fan environments, deposition of mud-rich sediments and air-fall tuffs within a lacustrine(?) environment, and subaqueous to subaerial extrusion of basaltic to dacitic arc volcanic rocks. This volcano-sedimentary section was intruded soon after deposition, with peak hornblende hornfels to low-pressure amphibolite facies metamorphism ultimately driven by intrusion of the surrounding Early Cretaceous Stokes Mountain ring dike complexes. Deposition of the nonmarine Goldstein Peak Formation within a fault-bounded, possibly transtensional, intra-arc basin during the transition from the low-standing, moderately extensional, Late Jurassic fringing arc to the high-standing, compressional mid- to Late Cretaceous arc indicates that at least one section of the Sierra Nevada arc was a fully emergent continental margin arc by the Early Cretaceous Epoch.
The Sierra Nevada range exposes one of the Cordillera’s most intensely studied batholiths. Enormous accumulations of granitoid magmas are the focus of studies directed at better understanding magmatic differentiation, magma emplacement, and the nature of continental growth at convergent margins (e.g., Kistler and Peterman, 1973; DePaolo, 1981; Saleeby, 1990; Glazner, 1991; Coleman et al., 1992; Sisson et al., 1996; Žák et al., 2009; Memeti et al., 2010a). A partial record of crustal growth events is also preserved in the fragments of metamorphic rocks that are stitched together by batholith magmas (e.g., Ernst, 1988; Saleeby, 2011). Integrative sedimentologic and tectonic studies of the Mesozoic Cordillera are directed at answering the question of how western North America evolved from the Paleozoic passive margin to the Late Cretaceous continental margin arc (e.g., Busby, 2004; Ingersoll, 2012).
The extensive late Cenozoic uplift, erosion, and glaciation that produced excellent exposures of the Sierra Nevada batholith also destroyed most records of depositional environments proximal to the Mesozoic arc. Paleoenvironmental reconstruction of the Mesozoic arc, therefore, relies on structurally isolated fragments of sedimentary and volcanic rocks that are preserved as roof pendants (e.g., Saleeby et al., 1978; Nokleberg, 1983; Schweickert and Lahren, 1991; Saleeby and Busby, 1993; Memeti et al., 2010b). Recent mapping of the Lake Kaweah pendant reveals the existence of a ∼2.4-km-thick metavolcano-sedimentary package that appears to be the youngest sedimentary record, as well as the only evidence of a nonmarine succession, within the framework of the Sierra Nevada batholith (Clemens-Knott, 2011). Following the guidelines of the North American Stratigraphic Code (Anonymous, 2005), these nonmarine metavolcano-sedimentary rocks are defined herein to compose the Goldstein Peak Formation.
In this paper we describe the metavolcano-sedimentary rocks and the metamorphic history of the Goldstein Peak Formation; reconstruct the protolith stratigraphy; and assemble the geologic history of deposition, foundering, intrusion, and syn-plutonic deformation. We interpret this newly recognized volcano-sedimentary succession in the context of the Early Cretaceous transition from a moderately extensional to a compressional arc that culminated with the emergence of the mid- to Late Cretaceous continental margin arc (Busby, 2004).
The Goldstein Peak Formation crops out on private agricultural and ranch land located in the westernmost foothills of the Sierra Nevada range, ∼35 km northeast of Visalia, California. The Goldstein Peak Formation forms the northwestern tip of the Lake Kaweah pendant and was previously included within the Late Triassic to mid-Jurassic Kings sequence, a marine metavolcano-sedimentary unit that composes the majority of the Lake Kaweah pendant (Fig. 1; Saleeby and Busby, 1993). The Goldstein Peak Formation was first recognized during detailed mapping of the informally designated Stokes Mountain region, a 325 km2 area of crystalline rocks noteworthy due to the unusually large and coherent exposures of mafic to intermediate plutonic rocks of the Early Cretaceous Sierra Nevada batholith (Saleeby and Sharp, 1980; Clemens-Knott, 2011).
The Stokes Mountain region is dominated by two ring dike complexes emplaced between 123 and 115 Ma (Fig. 2; Clemens-Knott and Saleeby, 1999). The Eastern and Western ring complexes expose the roots of Early Cretaceous arc stratovolcanoes that are separated by the syn-magmatic Stone Corral shear zone. This linear zone of intense magma mingling and limited left-lateral syn-magmatic displacement has been interpreted as some sort of local structure that accommodated crustal tumescence at 116 ± 2 Ma, during near-synchronous emplacement and activity in the overlying stratovolcanoes (Clemens-Knott and Saleeby, 1999). The ring dikes are dominated by pyroxene-hornblende-biotite tonalite to granodiorite; true granites are sparse.
The ring complexes are bordered to the east by the Lake Kaweah pendant, the westernmost outcrop of the Kings sequence at ∼36.5°N latitude (Fig. 1). The term “Kings sequence” was first used by Bateman and Clark (1974) for Upper Triassic and Lower Jurassic quartzites, carbonates, and locally interstratified metavolcanic rocks found between the Dinkey Creek and Mineral King regions. Saleeby et al. (1978) extended the usage to include similar metavolcano-sedimentary pendants distributed along the southwestern edge of the batholith as far south as the Tehachapi Mountains (∼35°N latitude). Those authors interpret the Kings sequence protoliths as being deposited in a shallow marine setting within an ignimbrite field, in which sedimentation and eruptive activity were contemporaneous. Ongoing mapping and research in the Stokes Mountain region is aimed at differentiating the Kings sequence from the Permian–Triassic Calaveras complex, which is composed of chert-argillite with associated carbonates (see discussion in Clemens-Knott and Saleeby, 2013). Mapping these two units is locally complex due to poor outcrop, overlapping rock types, and complex structural interrelationships. Understanding of the Kings sequence, the Calaveras complex, and other metasedimentary units is currently evolving, in large part due to detrital zircon investigations of the Western metamorphic belt and pendants throughout the Sierra Nevada batholith (e.g., Paterson, 2012; Clemens-Knott and Saleeby, 2013). For simplicity, in this paper the older Calaveras complex is not separated from the Kings sequence.
Roof pendants in the southern Sierra Nevada have been grouped into four terranes based on pre-metamorphic lithology. From west to east, the Kings-Kaweah, Kings, Goddard, and High Sierra terranes consist of ophiolite, proximal shallow marine sediments, volcanics, and marine sediments, respectively (Nokleberg, 1983; Saleeby and Busby, 1993). In this simplified organizational model, the Kings-Kaweah terrane is Ordovician–Permian to Middle Jurassic; the Kings and Goddard terranes were mostly deposited in the Mesozoic; and the High Sierra terrane is Precambrian to Paleozoic (Moore, 2000). The spatial distribution of terranes reflects a simplified reconstruction of the pre-plutonic geometry of the arc, wherein arc volcanic and sedimentary rocks are flanked by rocks of oceanic affinity to the west and rocks of passive-margin affinity to the east.
The Goldstein Peak Formation is sandwiched between rocks of the Kings-Kaweah ophiolite to the west and the marine Kings sequence to the east (Fig. 1). Intrusion of Early Cretaceous gabbros almost completely peeled the Kings sequence away from the Goldstein Peak Formation (Clemens-Knott, 2008), although fragments of the depositional unconformity between the Goldstein Peak Formation and the marine Kings sequence are preserved. Whereas the contact with the Kings sequence is partially preserved (Fig. 3; Clemens-Knott, 2011), the contact between the Goldstein Peak Formation and the older Kings-Kaweah ophiolite is not exposed. The location of this presumed tectonic contact is constrained by the Smith Mountain outcrops of pillow basalts of the Kaweah serpentinite mélange, located ∼18 km to the west of Goldstein Peak, and along the southwestern margin of the Eastern ring complex (Fig. 1).
The newly recognized Goldstein Peak Formation forms the northwestern section of the Lake Kaweah pendant adjacent to the Western ring complex (Fig. 2). The Goldstein Peak Formation is truncated to the south by the Eastern ring complex and is bisected obliquely by the Stone Corral shear zone. The Barton Mine, one of the southernmost Mother Lode hard-rock mines, is located where the shear zone cuts the northeastern border of the Goldstein Peak Formation (Anonymous, 1893, 1894, 1896). Mining focused on a ∼1-m-wide quartz vein associated with a biotite granite dike intruded into the Goldstein Peak Formation; elsewhere, the pendant is intruded by small dikes and pods of biotite-bearing to muscovite-garnet granite (Clemens-Knott and Saleeby, 1999).
Mapping was conducted at 1:24,000 scale (Clemens-Knott, 2011), and a ∼1250-m-thick stratigraphic section was measured using standard sedimentologic field techniques. Rock descriptions were augmented with petrographic analysis of standard thin sections (Table 1). Thermal ionization mass spectrometry (TIMS) U-Pb dating of bulk zircon samples (Tables 2A and 2B) was performed at the California Institute of Technology using methods similar to Saleeby et al. (1989) and Krogh (1973). Rocks were powdered in a tungsten carbide mill; whole-rock chemical analyses (Table 3) were performed by Activation Labs Ltd. using X-ray fluorescence (XRF) and inductively coupled plasma–optical emission spectrometry, and by XRF at Pomona College (see methods in Lackey et al., 2012).
Units of the Goldstein Peak Formation generally strike NNW, parallel to the elongate ridge that summits the 860-m-tall Goldstein Peak (36°37′44.82″N, 119°09′31.68″W). The near-vertical dip exposes a natural cross section through much of the formation (Fig. 3A, line A–A′), which is best exposed on a transect parallel to the northern bank of Rattlesnake Creek. Interbedded lenses of trough cross-stratified and low-angle planar cross-stratified metaquartzite within the quartz pebble metaconglomerates provide stratigraphic top indicators that point consistently to the west (Fig. 3B). The base of the Goldstein Peak Formation is preserved at its unconformable depositional contact with the Kings sequence; the original thickness of the formation is unknown, however, because the structural top of the formation is defined by the northwestern margin of the Lake Kaweah pendant (Fig. 2). Outcrop is generally poor in this never-glaciated and highly vegetated part of the Sierra Nevada foothills, so marker beds with atypical laterally continuous outcrop were mapped in an attempt to depict the formation’s internal structure (Fig. 3A). Interfingering between sedimentary protoliths is presumably more common than typically exposed.
Metamorphic Rock Types
Two types of metaconglomerates form the most prominent outcrops of the Goldstein Peak Formation: (1) quartz pebble metaconglomerates, interlayered with cross-bedded metasandstones, occur in the lower (eastern) two-thirds of the central section (Figs. 4A and 4B); and (2) polymict pebble metaconglomerates occur in the upper (western) third of the metamorphic section (Figs. 4C and 4D). Clast size ranges from pebbles to cobbles; pebble-sized clasts are predominant, with the abundance of cobble-sized clasts greatest in the polymict metaconglomerates. Most clasts preserve primary textures, but the matrix texture of the conglomerates is metamorphic.
The sillimanite-biotite-quartz schists are interbedded with several 10-cm- to 1-m-thick, quartzofeldspathic schists. Individual layers of quartz-rich schist display sharp boundaries with the surrounding biotite-rich schists, as well as relatively constant thickness along strike. Many of these resistant layers are laterally extensive (>1 km) and were mapped as marker beds within the biotite-rich schists (Fig. 3A).
Hornblende-rich amphibolites form a ∼3-km-long lens enclosed within the sillimanite-biotite-quartz schists (Fig. 3A). The lens thickness varies from 95 m in the north to 15 m in the south. The upper (western) contact is concordant with the overlying metasedimentary section; in the thickest, northern part of the amphibolite lens, the basal contact projects downwards into the stratigraphically lower metasediments. A unique, <1-m-thick zone containing 2- to 5-cm-long, bow tie– to star-shaped sillimanite clusters (Fig. 5A) commonly marks the basal (eastern) amphibolite-schist contact (at ∼250 m in Fig. 3B). Amphibolites display a variety of structural subtypes (Table 1). Most amphibolites have a globular structure in the thicker, northern part of the section; the dominant variety of amphibolite contains abundant ragged inclusions in the central section; and planar layering, with rare examples of cross-bedding, ripple marks, and channels, is found in the thin, southern section along with lenticular amphibolite.
Amphibolite whole-rock compositions display variations on Harker diagrams that are suggestive of liquid lines of descent within a magmatic system (Fig. 6). Interestingly, compositions of the quartzofeldspathic schist marker beds overlap the high-silica end of the amphibolite suite and seemingly extend the trends to even higher silica contents.
Relationship of the Goldstein Peak Formation to the Kings Sequence
Within the northern Lake Kaweah pendant, the Kings sequence is distinguished from the Goldstein Peak Formation by the presence of amphibolite-grade carbonate, calcsilicate, and marly amphibolite rocks. The Kings sequence is also recognized by pervasive deformation structures, such as isoclinal folds, open folds, and boudinage, indicative of possibly multiple, intense ductile deformation events (Saleeby and Busby, 1993). In the immediate vicinity of the contact between the Kings sequence and the Goldstein Peak Formation, the foliation in both units is roughly subparallel. The contact between the Goldstein Peak Formation and the Kings sequence, however, is typically obscured by poor outcrop and by multiple tabular tonalite-granodiorite intrusions along the southern end of the contact (Fig. 3A). A single exposure suggests an approximately conformable contact between marly amphibolite of the Kings sequence and net-veined biotite metaquartzite and pebble metaconglomerate of the Goldstein Peak Formation. A thin, poorly exposed layer of schist of indeterminate affinity separates the marl, presumably of the Kings sequence, from quartzite, presumably of the Goldstein Peak Formation. Stratigraphic-top indicators consistently point westward indicating that the Goldstein Peak Formation was deposited on top of the Kings sequence. This relationship requires that the Goldstein Peak Formation be younger than the Jurassic–Triassic Kings sequence, a relationship consistent with the higher metamorphic grade and more extreme deformation history of the Kings sequence (Saleeby and Busby, 1993).
Previous U-Pb geochronologic investigations of the Stokes Mountain region constrain the age of ring-dike plutonism to between 123 and 115 Ma (Saleeby and Sharp, 1980; Clemens-Knott and Saleeby, 1999). We note that an earlier published date of 113 ± 3 Ma (Chen and Moore, 1982) from the Eastern ring complex has been revised to 115 Ma (Lackey et al., 2005), consistent with our preferred span of ring-dike plutonism. In this study, three concordant U-Pb zircon TIMS dates provide new information regarding the geologic history of the Stokes Mountain region (Tables 2A and 2B):
1) Late Jurassic skarn formation in the Kings sequence: The Consolidated Tungsten Mine skarn is located on the northeastern tip of the Lake Kaweah pendant (Fig. 3A), ∼3 km from the contact with the Goldstein Peak Formation. Mineral associations in the skarn indicate amphibolite-facies metamorphic conditions (Newberry, 1980; Berekian and Clemens-Knott, 2006). Dating of the associated quartz diorite pluton (sample MM100-2B) indicates that skarn formation occurred in the late Middle Jurassic (164 ± 1 Ma). This pluton is coeval with the ca. 170–156 Ma Mill Creek complex, which crops out just 15 km to the northwest (Saleeby, 2011) and is part of a longer belt of Jurassic ultramafic to intermediate plutonic rocks that extends through the northern Sierra Nevada and into the Klamath Mountains (Saleeby and Sharp, 1980; Snoke et al., 1982). This date provides a minimum depositional age constraint for the Kings sequence of the northern Lake Kaweah pendant (Fig. 2) as well as a maximum age constraint for the overlying Goldstein Peak Formation.
2) Early Cretaceous wet-sediment intrusion: A swarm of garnet-bearing granite dikes intrudes metaconglomerates that overlie the amphibolite lens. Dike contacts display cuspate margins and irregular, curvilinear apophyses that are strongly suggestive of intrusion into wet sediment (Fig. 4F). This structural relationship implies that the Early Cretaceous age of the granite dike (139 ± 1 Ma; sample SM286) places a minimum constraint on the depositional age of the Goldstein Peak Formation.
3) Mid-Cretaceous termination of ring-dike intrusion: The Lake Kaweah pendant structurally defines the eastern and northern boundaries of the Western and Eastern ring complexes, respectively, apparently separating the ring complexes from an adjacent magmatic center located to the north. Magmatic foliation and enclave swarm trends within this adjacent, hornblende-biotite tonalite are roughly concordant with the east-west–trending Kings sequence ridge that forms the northern margin of the Lake Kaweah pendant (Fig. 3A). Our preferred interpretation is that this 114 ± 1 Ma pluton (sample TM124) truncates the Stone Corral shear zone and provides a minimum age boundary for the termination of Stokes Mountain ring-dike plutonism. This age is indistinguishable from two Ar-Ar cooling ages from layered cumulate megaxenoliths that are engulfed by the ring complexes (Slopko et al., 2003), suggesting that the Stokes Mountain region cooled quickly after the ca. 115 Ma completion of ring-dike emplacement, with the focus of magmatism shifting further to the north and east (Saleeby and Sharp, 1980; Lackey et al., 2012).
Protoliths of the Goldstein Peak Formation
Based on preserved sedimentary and volcanic structures, as well as mineral and whole-rock compositions, the protoliths of the Goldstein Peak Formation are clearly sedimentary and volcanic (Fig. 3B). Moreover, in all but the schist unit, the limited degree of metamorphic recrystallization permits fairly specific protolith interpretations and supports a detailed reconstruction of Early Cretaceous depositional environments (Fig. 3C).
The two major conglomerate lithofacies likely were deposited by different mechanisms. Deposition of the typically clast-supported quartz pebble conglomerates by traction currents in moving water is indicated by the association of cross-bedded conglomerates, cross-laminated sandstones, and imbricated pebbly sandstones (Miall, 1977). In contrast, the polymict pebble conglomerates commonly are matrix supported and lack internal organization (e.g., cross-bedding, imbrication), indicating deposition of this lithofacies in a regime dominated by sediment-gravity flows (Lowe, 1979). The greater abundance of biotite in the biotite-quartz matrix of the polymict conglomerates is consistent with a significant original clay component; moreover, 2–5-cm-long, curved lenses of biotite found in some horizons resemble mud intraclasts. Interlayered with the polymict metaconglomerates are biotite- and pebble-poor beds that do not, however, have the characteristics of sediment-gravity flow deposits. For example, one package consists of a 40-cm-thick gravel-rich layer sandwiched between two, 20-cm-thick beds containing fine to medium pebbles in a fine to coarse sandy matrix (Fig. 4E). The lower abundance of biotite, better sorting, and packaging pattern of these three conglomerate beds are consistent with deposition of the pebbly sandstones by rapid water-flow depositional mechanisms (i.e., sheet flood; Bull, 1972; Hogg, 1982).
Deposition of the polymict pebble conglomerates by gravity flow punctuated by sheet flood events is consistent with an alluvial fan environment (Bull, 1972; Blair and McPherson, 1994). The significant range of size and angularity of clasts, along with the generally greater matrix abundance, support this interpretation (Figs. 3B and 3C). In contrast, water-laid traction is consistent with deposition of the quartz pebble metaconglomerates in a fluvial environment (Miall, 1977). Cross-bedded conglomerates suggest flow rates and bed loads consistent with braided fluvial systems (Khadkikar, 1999).
Relatively long transport distances for the quartz pebble conglomerates are suggested by the high degree of rounding and good textural, size, and compositional sorting of the clast population (Sneed and Folk, 1958; Mill, 1979). In contrast, much of the diverse pebble population in the polymict pebble metaconglomerates could have been derived locally. Specifically, clasts of the texturally distinctive amphibolites, schists, and biotite quartzites suggest reworking of lower stratigraphic horizons of the Goldstein Peak Formation. Similarly, quartz pebbles, particularly compound clasts composed of fine-grained quartzite bounding coarse-grained quartz pebble fragments, may have been derived from the older quartz pebble conglomerates of the Goldstein Peak Formation. Additional metamorphic sources for the clasts are not required, though dark chert clasts found in conglomerates immediately underlying the amphibolite lens may have been derived from the underlying Kings sequence or Calaveras complex (Saleeby and Busby, 1993). Abundant clasts of hornblende-biotite granitoids, along with a single anorthositic gabbro clast, suggest some amount of exhumation and erosion of Jurassic–Triassic arc plutons prior to conglomerate deposition. Lastly, elongate clasts of fine-grained mafic and felsic rock types are typically deformed around adjacent quartz pebbles and plutonic clasts (Fig. 4D). The implied weakness of this bimodal population of fine-grained lenses, as well as the relict porphyritic textures, are consistent with a young volcanic origin. Taken together, clasts in the polymict metaconglomerates may have been derived from the Jurassic–Triassic arc complex, as well as from underlying units in the Goldstein Peak Formation, and possibly from the Early Cretaceous volcanic arc. The contrast between local clast derivation in the polymict pebble conglomerates and distant clast derivation for the quartz pebble conglomerates is consistent with a transition from long-system fluvial deposition of the lower Goldstein Peak Formation to short-system alluvial fan deposition of the uppermost Goldstein Peak Formation.
Extensive metamorphic recrystallization has obliterated the protolith mineralogy and any original depositional structures in the schists. The abundance of biotite indicates high aluminum and hydroxide components, which together suggest a large original clay content (Fedo, 2000). Fine-grained particles, such as silt and clay, are indicative of lower-energy environments. Given the association of schists with the fluvial and alluvial fan conglomerates, likely depositional environments of the pelite protoliths include lacustrine, overbank, or estuarine to shallow marine. However, without any fossils or knowledge regarding the geographic extent of this unit, and given the fact that alluvial fan and braided stream deposits could prograde into any of these environments, the depositional environment(s) of the schists remain unclear.
Thin calcsilicate-bearing quartzofeldspathic marker beds are mapped in the schists overlying the amphibolite lens (Fig. 3A). The whole-rock compositions of these marker beds are grossly rhyolitic and plot along chemical trends defined by the mafic to intermediate amphibolites (Fig. 6). Volcanic glass would be highly reactive in water, possibly accounting for the extensive recrystallization of these high-silica rocks relative to the quartz-rich clastic metasediments found elsewhere in the section; moreover, some amount of chemical alteration would be expected during both lithification and metamorphic recrystallization. If the protoliths of these laterally extensive marker beds contained large components of silicic ash, then we favor deposition of the surrounding sillimanite-biotite-quartz schists in a relatively low-energy lacustrine environment (Fig. 3C; Smith and Landis, 1995).
Whole-rock compositions of the typical amphibolites range from basalt to low-silica dacite (Table 3). Relatively smooth geochemical variations are consistent with overall preservation of igneous compositions during extrusion and contact metamorphism (Fig. 6), although some mobility of Ti and P is indicated. Metamorphic recrystallization appears to have obliterated the original grain-to-grain contacts but has only partially destroyed the larger phenocrysts and glomerocrysts of pyroxene and plagioclase. Trace element abundances of a massive xenolith-bearing amphibolite (sample DS-2b, Table 3) sampled from the base of the thick, northern end of the section classify this sample as an island-arc basalt of transitional calc-alkaline to tholeiitic affinity (Pearce, 1982). Its high Mg# (74; Table 3) suggests a mantle derivation for the volcanic protolith (e.g., Winter, 2010).
The composition of sample DS-4 (Table 3), however, is more similar to the marker beds found in the overlying schists than to the typical amphibolites. A pod of this white mica-sillimanite-quartz schist (Table 1) is found near the base of the amphibolite lens. This schist is both mineralogically and texturally similar to the stratigraphically overlying quartzofeldspathic marker beds. We speculate that the protoliths of the rare white schist (DS-4), as well as the associated “nubby” lithofacies (Table 1), are high-silica ash and pumice deposits. Protolith interpretations for the various textural varieties of amphibolite (Table 1, Fig. 5) are presented in approximate geographic order, from north to south.
Globular amphibolite. Particularly distinctive volcanic rocks called hyaloclastites and peperites are present in a number of Sierra Nevada pendants (Hanson and Schweickert, 1982; Saleeby et al., 1978). The term hyaloclastite refers to metavolcanic breccias formed by non-explosive fracturing and disintegration of lavas and intrusions quenched by water (see discussion in McPhie et al., 1993). A peperite is a hyaloclastite generated by the mixing of coherent lava or magma with unconsolidated sediment. The degree of sorting of the sediment into which magma intrudes will control the three-dimensional shape of inclusions, thereby changing the type of peperite formation (i.e., fluidal “globular” peperite versus blocky peperites; Busby-Spera and White, 1987).
Clast shapes and the geometry of monolithologic clast swarms suggest that the protoliths of the lenticular amphibolites were hyaloclastites with lesser peperites. Because the presence of amphibole appears to reflect an original mafic volcanic component, the amphibole-free zones both surrounding clasts and forming lenses are interpreted as intercalated siliciclastic sediment. Rounded quartz grains found within the pillow matrix also reflect the original sedimentary component (Fig. 5B). The protolith of the massive, xenolith-bearing amphibolite (sample DS-2b, Table 1), located at the base of the amphibolite section (at ∼250 m in Fig. 3B), is interpreted as either a lava dome or flow, with the olive-green, concordant tabular body higher up in the section originating as either a lava flow or sill (Table 1).
Ragged amphibolite. Protoliths of the ragged facies are interpreted as being tephra deposits that contained ragged pumice lapilli and in which biotite and quartz reflect an original sedimentary component. Orange quartzofeldspathic spheres found within this unit may have originated as accretionary lapilli. This intermediate zone contains some horizons of globular amphibolite (dominant to the north) and some bedded amphibolite (dominant to the south).
Lenticular and layered amphibolites. The amphibolite unit thins and ultimately pinches out to the south where it is directly underlain by quartz pebble metaconglomerate. This thin, southern end of the amphibolite lens is composed of 2–5-m-thick, unsorted medium to light gray amphibolite horizons that contain abundant lensoidal porphyritic clasts (Fig. 5E). Some of these lenticular horizons are truncated by channels having apparent widths up to 15 m. Above the lenticular amphibolites are packages of layered amphibolites (Fig. 5F). Compositional layers are commonly oblique to both regional and mineral foliations, and are interpreted as preserved bedding. As such, the southern section is interpreted as having formed from repeated extrusion of ash flows across a fluvial braid plain, in which the channels, cross-bedding, and laminae indicate fluvial reworking of the volcanic deposits.
The mapped variation from a thicker northern amphibolite section, with subaqueously extruded volcanics, to the thinner southern amphibolite section, with reworked volcanics (Fig. 3A), suggests a facies change controlled by water depth: a depositional basin centered in the northern section, grading southward across a broad shoreline (in modern directions). The sillimanite-star zone that commonly marks the base of the amphibolite section (Fig. 5A) is interpreted as a recrystallized contact metamorphic zone, initially formed in a mudstone protolith overlain by hot volcanic rocks.
Mineralogical and textural reequilibration of the quartz-rich and mafic protoliths was limited. Mineral stability relations in the schist constrain the metamorphic grade, and the extent of deformation can be evaluated in the pebble metaconglomerates, wherein the quartzite pebbles resisted most deformation.
Metamorphic grade is estimated as hornblende hornfels to low-pressure amphibolite facies, based on the incomplete replacement of andalusite by sillimanite (var. fibrolite) and by the presence of cordierite. Stability relations indicate metamorphic pressures less than ∼3 kbar and temperatures of ∼600–700 °C (Powell and Holland, 1990; Powell et al., 1998; Winter, 2010). Relatively low pressures are confirmed by reaction relationships observed between white mica, andalusite, cordierite, and biotite, though higher pressures may be indicated by rare staurolite. The onset of partial melting of biotite-bearing rocks is indicated by the common occurrence of net-veining within the biotite quartzites at the pendant margins, as well as the presence of small dikes and pods of tourmaline-garnet granite. Such hot, upper-crustal conditions are consistent with field relations indicating that the septum separates two Early Cretaceous ring dike complexes that enclose stoped xenoliths of coeval hypabyssal rocks (Clemens-Knott and Saleeby, 1999). Retrograde recrystallization is minor, limited to partial replacement of biotite by chlorite in the schists, as well as partial replacement of hornblende by biotite and saussuritization of plagioclase feldspar in the amphibolites.
Assessing the effects of metamorphism in the amphibolites is difficult because the primary igneous minerals (e.g., plagioclase, hornblende, quartz) are all stable within the hornblende hornfels to amphibolite facies. Partial to complete recrystallization is indicated by the development of mosaic textures in the originally fine-grained sectors of the rocks, and by the inferred replacement of sedimentary clay minerals by biotite. Metamorphic minerals such as sillimanite, andalusite, tremolite-actinolite, and cummingtonite-grunerite also record metamorphic recrystallization, and certain minerals (e.g., calcite, serpentine, quartz veins) suggest limited degrees of fluid-aided retrograde recrystallization.
Observations that suggest a short-lived metamorphic event include the outcrop-scale preservation of many primary structures (e.g., cross-bedding, and ragged and lenticular clast margins); the incomplete recrystallization of volcanic textures (e.g., large sieve-textured plagioclase phenocrysts and plagioclase glomerocrysts); and incomplete prograde metamorphic reactions (e.g., andalusite to sillimanite). Pebbles and cobbles display variable, but generally limited, degrees of textural reequilibration, consistent with the preservation of sedimentary and volcanic structures.
At multiple locations throughout the Goldstein Peak Formation, metasediments are crosscut by groups of regularly spaced, tube-like structures that are ellipsoidal in cross section and have maximum diameters of 7–20 cm. In the metaconglomerates, tube boundaries cut across individual pebbles indicating a post-deposition origin. Many tubes display a concentric color zonation (Fig. 7A); in the tube cores, epidote and quartz partially replace the protolith mineralogy while preserving the clastic texture of the protolith. A preliminary interpretation of these structures is that they represent fluid-expulsion tubes, possibly resulting from rapid heating fueled by wet-sediment intrusion (Fig. 4F). Clearly, further analysis is necessary to test this hypothesis, but it implies that metamorphism may have begun prior to complete lithification of the volcano-sedimentary section.
Deformation of the Goldstein Peak Formation
Within the Goldstein Peak Formation, map-scale foliation is everywhere subparallel to the compositional stratification. The metamorphic foliation is defined by biotite and sillimanite, as well as by the primary and secondary axes of quartz pebbles. On outcrop scale, foliation of the fine-grained biotite-quartz matrix wraps around pebbles and cobbles, and quartz-rich areas typically display mosaic textures. Compaction of the metaconglomerates is demonstrated by the deformation of schist and metavolcanic clasts around quartzite clasts (Fig. 4D). On the basis of the geometry of cross-beds and the dimensions of globular amphibolites, there appears to be no other deformation than that expected for the presumed depth of burial. There is limited evidence for focused ductile deformation (e.g., undulatory quartz, bent albite twins) and brittle deformation (e.g., biotite and sillimanite filling fractures in quartzite pebbles).
Zones of intense ductile deformation obliquely cut the metamorphic foliation. North-trending shear zones are identified by mylonitic fabrics that are characterized by quartz ribbons as well as augen of andalusite, white mica, and biotite. Where intersecting pebble metaconglomerates, the atypically stretched clasts exceed dimensions of 1:10, and in one outcrop, folded cobbles indicate significant amounts of focused strain (Figs. 7B and 7C). The poor lateral continuity of outcrop prohibits tracing individual shear zones for much distance, but foliations suggest that these zones of focused ductile deformation are subparallel to the NNE-trending Stone Corral shear zone (Fig. 2). The Stone Corral shear zone was originally interpreted as a syn-magmatic crustal tear that accommodated crustal tumescence in the adjacent, coeval ring dike complexes (Clemens-Knott and Saleeby, 1999). Observed left-lateral shear was explained by the resolution of ring-complex radial expansion along the trace of the shear zone (Fig. 8).
Though the Stone Corral shear zone accommodated a small amount of left-lateral shear, the dominant structural character appears to have been the accommodation of focused magma intrusion and mingling of a variety of magma types where the ring complexes abut each other (Clemens-Knott and Saleeby, 1999). Isolated mylonitic outcrops document projection of the shear zone to the north through the plutonic rocks that are in contact with the Goldstein Peak Formation, but there is no indication that the shear zone bisects the east-west–trending ridge of Kings sequence that forms the northern boundary of the Lake Kaweah pendant (Fig. 3A). Immediately west of a foliated pyroxene gabbro that marks the northernmost documented location of the Stone Corral shear zone (point A in Fig. 8), a single outcrop of Goldstein Peak schist preserves a second foliation. Here, a west-vergent foliation partially transposes the dominantly NNW-trending, near-vertical foliation of the Goldstein Peak Formation (Fig. 7D; point B in Fig. 8).
The southern half of the Goldstein Peak Formation is almost completely surrounded by Early Cretaceous tonalitic to granodioritic plutons. Although thin, discontinuous strips of the Kings sequence remain connected to the base of the Goldstein Peak section, the majority of the Kings sequence appears to be removed, replaced by a concentration of hornblende- and pyroxene-rich gabbros that is centered roughly on the community of Auckland (Fig. 2). Enclosed within the northwestern sector of the Auckland gabbros are an abundance of meter- to decameter-long xenoliths of Kings sequence (Fig. 3A). Xenolith foliations suggest a progressive counterclockwise rotation of the Kings sequence, from a NNW orientation parallel to the trend of the Goldstein Peak Formation, to an east-west orientation. These field observations suggest that the Kings sequence was dismembered and rotated during forceful intrusion of the Auckland gabbros (Fig. 8). Moreover, the rotation that made space for gabbro intrusion was in part accommodated by left-lateral shear along the northern end of the Stone Corral shear zone. The enigmatic west-vergent foliation noted above (Fig. 7D; point B in Fig. 8) is located in what would be the “hinge zone” that would have accommodated the ∼60° counterclockwise rotation of dismembered Kings sequence that now forms the prominent east-west–trending ridge. Emplacement of the water-rich Auckland gabbros may record the formation of a satellite cone on the northern flank of the stratovolcano that capped the Eastern ring complex. The curving map pattern of the northern Stone Corral shear zone suggests its origin as a ring fault bounding the southeastern sector of the Western ring complex (Fig. 8).
Geologic History of the Goldstein Peak Formation
Fluvial sediments composing the exposed base of the Goldstein Peak Formation were deposited on top of the previously deformed and metamorphosed Kings sequence. Locally, the age of the Kings sequence is constrained to earlier than latest Middle Jurassic (164 ± 1 Ma) by the Consolidated Tungsten Mine skarn pluton (Table 2A). Moreover, the absence of a penetrative deformation in the Goldstein Peak Formation, which is attributed elsewhere in the region to the ca. 150 ± 2 Ma “Nevadan” orogeny (Wolf and Saleeby, 1992), supports a post–150 Ma initiation of Goldstein Peak Formation deposition.
Assuming an entirely depositional contact (Fig. 3A), the base of the Goldstein Peak Formation appears to lap to the north (modern direction) upsection across the Kings sequence, implying that this part of the Kings sequence formed a local topographic high (Fig. 9A). Deposition of well-rounded, quartz-rich, fluvial conglomerates and sandstones was succeeded by deposition of sandy mudstones (Figs. 3A and 3B). Complete metamorphic recrystallization of the mud-rich sediments has destroyed any fossil evidence that might constrain the depositional environment as marine (e.g., estuarine, shallow marine) or nonmarine (e.g., overbank, lacustrine), but the apparent rapid transition from fluvial sediments and, upsection, back to fluvial, supports our preferred interpretation of a nonmarine depositional environment for the mudstones (Fig. 3C).
Sediment deposition in a low-energy aqueous environment was locally interrupted by the extrusion of mafic magma. The map pattern (Fig. 3A), coupled with the north-to-south variation in volcanic textures described above, suggests that volcanism centered on a sub-basin that was deepest in the north (modern direction) and shallowed to the south. Globular peperites and hyaloclastites, possibly with rare lava flows, were extruded in the sub-basin; to the south, volcanic ash flows interfingered with fluvial sediments. Fluvial reworking of volcanic deposits produced the layered and cross-bedded amphibolites found near the thin southern end of the amphibolite lens. Deposition of muddy to sandy sediments continued to both the north and south (modern directions), reconnecting after the termination of mafic volcanism to form what ultimately constituted a >250-m-thick blanket of mud-rich sediments over the volcanic sub-basin. Thin, laterally extensive, quartzite marker beds within the sillimanite-biotite-quartz schists (Fig. 3) are interpreted as volcanic ash beds preserved within the section of sandy mudstones overlying the amphibolite lens (Figs. 3A–3C). If this interpretation is correct, proximal subaqueous mafic volcanism was succeeded by distal, explosive silicic arc volcanism.
Upsection framework-supported quartz pebble conglomerates, with lenses of cross-laminated sandstones and rare cross-bedded conglomerates, record the return of fluvial deposition. These lithofacies, however, are replaced farther upsection by matrix-supported, polymict pebble conglomerates containing lenses of massive, coarse sandstones that together may represent alluvial fan deposits (Figs. 3B and 3C). Poorly sorted clasts, having a variety of compositions, demonstrate reworking of underlying Goldstein Peak lithofacies (e.g., distinctive amphibolites, well-rounded quartz pebble clasts) and suggest the beginning of dissection of the Jurassic–Triassic arc (e.g., granitoid with lesser gabbroid clasts). Felsic clasts that have relict porphyritic textures and are typically deformed around crystalline clasts may record the erosion of weak, silicic volcanic rocks (Fig. 4D). The presumed change in depositional facies, combined with the presence of a bimodal population of volcanic clasts, arc plutonic clasts, and clasts derived from underlying lithofacies, are together consistent with normal faulting and basin formation within an extensional arc environment.
A swarm of ∼0.5-m-thick garnet-biotite granite and biotite granite dikes intrudes the clastic section overlying the amphibolites. Cuspate dike margins and curvilinear apophyses suggest that the dikes intruded prior to complete lithification of the sediments (Fig. 4F). Clusters of quartz-epidote tubes that crosscut sedimentary clasts are speculated to be water-expulsion pathways formed in rapidly heated, wet sediments (Fig. 7A). The 139 ± 1 Ma age of wet-sediment intrusion (Table 2A) seemingly constrains the post-“Nevadan” deposition of the Goldstein Peak Formation to between ca. 150 and ca. 138 Ma.
No record of near-surface events occurring between ca. 138 and 126 Ma has yet been recognized within the immediate Stokes Mountain region, though burial and rotation of the Goldstein Peak Formation must have occurred during this time span (Fig. 9B). Based on the previously discussed evidence for sedimentation within normal fault–bounded basins (Fig. 3C), we hypothesize that much of the rotation of the northern Lake Kaweah pendant occurred by listric normal faulting in an extensional regime in a manner similar to that proposed for the Early Cretaceous Goddard and Ritter volcanic sections (Tobisch et el., 1986). At depth during this period, however, the ring-dike magmatic system was beginning to develop: the oldest 13% of 261 individually dated zircons separated from Stokes Mountain region gabbroids to granitoids crystallized between 135 and 126 Ma (Gevedon and Clemens-Knott, 2012).
Sometime in the middle Early Cretaceous (i.e., Aptian), mafic magma intruded the Stokes Mountain region once again, this time forming a suite of layered olivine-plagioclase gabbros (Fig. 9C; Clemens-Knott and Saleeby, 1999). At present, the age of these cumulates is constrained only by a single K-Ar date of 123 ± 3 Ma (Saleeby and Sharp, 1980). Stratigraphic variation of olivine and plagioclase compositions within the layered gabbros documents repeated recharge of primitive mafic magmas, which suggests the possibility of coeval mafic volcanism (Seal, 2011; Clemens-Knott et al., 2011). Beginning at ca. 123 Ma, mafic to felsic plutonism formed the Western and Eastern ring complexes, which have an average composition of tonalite (Fig. 2; Clemens-Knott and Saleeby, 1999). Volcanism accompanying ring-dike emplacement is supported by a 120 ± 1 Ma concordant date of a low-δ18O, hypabyssal-textured xenolith surrounded by a 120 ± 1 Ma, normal-δ18O tonalite (Clemens-Knott and Saleeby, 1999) that documents stoping of the hypabyssal-volcanic carapace by the underlying ring dikes. As these subvolcanic centers inflated and began to impinge upon each other at ca. 116 ± 2 Ma (Clemens-Knott and Saleeby, 1999), crustal tearing was accommodated by a limited amount of left-lateral shear along the NNE-trending Stone Corral shear zone ring fault (Fig. 8). At the northern terminus of the shear zone, left-lateral displacement sheared the Kings sequence away from the Goldstein Peak Formation. Coupled with forceful emplacement of the Auckland gabbros (Clemens-Knott, 2008), a section of Kings sequence was “peeled away” from the Goldstein Peak Formation and rotated ∼60° counterclockwise to form the prominent east-west–trending ridge that forms the northern boundary of the Lake Kaweah pendant. Counterclockwise rotation during plutonism is recorded by the range of NNW to east-west foliations of Kings sequence xenoliths preserved within the Auckland gabbros and by the partial top-to-the-west transposition of Goldstein Peak foliation in the “hinge zone” (Fig. 7D).
This Early Cretaceous magmatism drove the metamorphic recrystallization of the Goldstein Peak Formation under conditions of hornblende hornfels to lower amphibolite facies. Metamorphism is estimated to have occurred at shallow crustal depths (Clemens-Knott and Saleeby, 1999). Temperatures high enough for partial melting of the Goldstein Peak Formation were reached at pluton-pendant contacts, as evidenced by the common occurrence of biotite quartzites cut by granitic net veins. Reconnaissance study of the polymetamorphic Kings sequence within the northern Lake Kaweah pendant, however, has not revealed a clear recrystallization overprint attributable to the Early Cretaceous metamorphic event (Castellanos and Clemens-Knott, 2010).
Ar-Ar data from layered gabbros suggests that the Stokes Mountain region cooled beneath the closure temperature of hornblende by 112.4 ± 1.2 Ma (Slopko et al., 2003). Since then the region, which is currently exposed at elevations of ∼120–975 m above sea level, experienced a net uplift of ∼4–6 km, and an unknown thickness of volcanic and sedimentary rocks was removed by erosion. Minimal Cenozoic uplift of the Stokes Mountain region is supported regionally by its location on the Sierra Nevada “hinge zone” of Wahrhaftig (1962). The “buried” topographic signature of the region is interpreted as indicating a recent depression and partial burial. Saleeby and Foster (2004) interpreted this sinking as resulting from the lithospheric downwarping driven by downwelling and detachment of the mantle drip that generated the Lake Isabella geophysical anomaly.
Tectonic Setting During Deposition of the Goldstein Peak Formation
Protoliths of the metamorphic Goldstein Peak Formation consist of fluvial, lacustrine, and alluvial fan sediments and mafic to intermediate volcanic and volcaniclastic rocks that formed coincident with the earliest stages of the Cretaceous Sierra Nevada batholith. The nonmarine origin of at least part of this Early Cretaceous section distinguishes the Goldstein Peak Formation from all other Sierra Nevada pendants and from virtually all other coeval deposits associated with the Sierra Nevada arc. Identifying the local and regional tectonic environment in which the Goldstein Peak Formation was deposited, therefore, illuminates a poorly understood period in the evolution of one of the classic ancient arc–forearc–accretionary wedge systems.
Upturned marine sediments exposed along the western edges of the Sacramento and San Joaquin Valleys record the latest Jurassic to earliest Cretaceous stabilization of the Great Valley forearc basin (Ingersoll and Dickinson, 1981; Suchecki, 1984; Dickinson, 1995; Surpless et al., 2006). Though the eastward subsurface extent of the Early Cretaceous Great Valley Group is unknown, interpretive cross sections across the forearc axis imply termination of Early Cretaceous deposits far west of the Stokes Mountain region. The position of the earliest Cretaceous arc is also poorly constrained, as virtually the entire arc is thought to be buried beneath the eastern half of the modern Great Valley (Clemens-Knott and Saleeby, 2013); rocks of the Stokes Mountain region, along with the Fine Gold Intrusive Suite just to the north, constitute the major exposures of plutonic rocks formed during the latter stages of the Early Cretaceous batholith (e.g., 123–115 Ma; Clemens-Knott and Saleeby, 1999; Lackey et al., 2012). Given these constraints, it is reasonable to conclude that the Goldstein Peak basin was located east of the marine forearc basin.
An Early Cretaceous backarc basin setting for the Goldstein Peak basin is unlikely, given the lack of evidence documenting marine deposition or formation of oceanic crust associated with rifting during this period. Moreover, the voluminous, texturally and compositionally mature, quartz pebble conglomerates that dominate the lower Goldstein Peak section imply derivation from highlands to the east, originating perhaps from exhumed Triassic–Jurassic arc plutonic rocks or from erosion of craton-derived quartz-rich sediments previously deposited in the arc’s central graben depression (Busby-Spera, 1988). Taken together, the most reasonable conclusion is that the Goldstein Peak Formation was deposited in an intra-arc basin that formed on the Triassic–Jurassic Sierra Nevada arc massif and within the nascent Cretaceous continental margin arc.
Intra-arc basins form in low areas bounded by volcanic edifices, in fault-bounded depressions, or in some combination of these structures (Smith and Landis, 1995; Miall, 2000). Though reconstructing the extent of the earliest Cretaceous volcanic edifice is impossible, the appearance of alluvial fan deposits in the upper third of the formation is consistent with at least partial fault bounding of the Goldstein Peak basin. The predominance of nonvolcanic, distal clastic deposits in the lower part of the Goldstein Peak section suggests that the basin was located far from active volcanic edifices (Smith and Landis, 1995). Moreover, deposition within a half graben, bounded on the west by a normal fault, provides a likely mechanism for producing significant down-to-the-west listric normal rotation of the section (Fig. 9B) prior to its being engulfed ∼15–20 m.y. later as the arc migrated into the Stokes Mountain region. Deposition of the Goldstein Peak Formation in an extensional regime would be consistent with the multiple traces of coeval bimodal volcanism: the lens of mafic volcanic rocks; the quartzofeldspathic marker beds within the schists, which are interpreted as former silicic ash beds preserved in lacustrine muds; and the apparently bimodal population of volcanic clasts within the alluvial fan deposits.
Considered in the context of Busby’s (2004) model for the evolution of convergent margins, the Goldstein Peak Formation constitutes a rare depositional record of the end of the mild extensional stage affecting the central Sierra Nevada arc (i.e., phase 2; Busby, 2012). Independent evidence for Early Cretaceous extension prior to ca. 123 Ma includes a change from nonaccretionary to accretionary behavior recorded in the Franciscan Complex (Dumitru et al., 2010) and subsurface structures in the forearc basin that may be interpreted as recording Early Cretaceous syn-depositional extension (Constenius et al., 2000). A switch from mildly extensional to compressional strain at 123 Ma locally coincides with the initiation of mafic to intermediate ring-dike magmatism in the Stokes Mountain region and documents the earliest stages of the eastward migration of the Cretaceous arc axis.
Early Cretaceous Arcs of California
The mid- to Late Cretaceous Sierra Nevada and Peninsular Ranges arcs of California are widely accepted as representing high-standing continental margin arcs. The nature of these arcs during the Early Cretaceous is, however, less well known and likely varied along strike. For example, in the Peninsular Ranges, marine strata inboard of the Early Cretaceous Alisitos arc indicates this was an oceanic arc, though its specific origin either as a fringing arc located in front of a narrow, backarc basin, or as an exotic arc, is debated (Wetmore et al., 2002; Busby, 2004). The elevation of the Peninsular Ranges arc varied along strike, with at least part being above sea level (the El Rosario segment; Fackler-Adams and Busby, 1998). The apparent absence of an accretionary wedge (Sedlock, 1996) and the lack of Early Cretaceous forearc sedimentation associated with the Peninsular Ranges arc contrasts with the Late Jurassic to Early Cretaceous stabilization of the Great Valley forearc basin (Ingersoll, 1982; Surpless et al., 2006). Given these differences in the forearc regions, it should not be surprising that the Early Cretaceous arcs themselves might also have differed, from oceanic in the Peninsular Ranges arc to continental in the central Sierra Nevada arc.
The newly defined Goldstein Peak Formation preserves a rare landscape record documenting the emergence of the Cretaceous Sierra Nevada continental margin arc. The dominantly siliciclastic formation was deposited in an intra-arc basin located slightly east of the initial axis of the Early Cretaceous arc and formed within the Triassic–Jurassic Sierra Nevada arc massif. Fluvial systems deposited clasts derived from exhumed arc plutonic rocks and associated quartz-rich sediments; associated mud-rich sediments record overbank or lacustrine facies. This fluvial-lacustrine environment was temporarily interrupted by the subaqueous eruption of mafic to intermediate arc volcanic rocks within a sub-basin; later still, the fluvial-lacustrine environment was succeeded by deposition within alluvial fans formed on the margin of the fault-bounded intra-arc basin. Continued listric normal faulting rotated the Goldstein Peak Formation to a near-vertical orientation. At ca. 123 Ma, the local strain regime transitioned from mildly extensional to compressional. The local expression of this transition was the intrusion of sub-volcanic Stokes Mountain ring dike complexes that engulfed, metamorphosed, and deformed the Goldstein Peak Formation. With time, the axis of Cretaceous arc magmatism migrated farther eastward, forming the high-standing mid- to Late Cretaceous arc, while younger forearc deposits buried the majority of the Early Cretaceous arc.
This paper was significantly improved by the thoughtful reviews of Jade Star Lackey and Kathy Surpless, as well as by Cathy Busby’s editorial handling. This study was supported by student-faculty research grants from the California State University (CSU) and by National Science Foundation (NSF) grant EAR-9105692 to Saleeby. Initial descriptions of the metasedimentary and metavolcanic sections constituted the undergraduate thesis projects of CSU Fullerton (CSUF) alumni Dolores van der Kolk (Ph.D., in progress, University of Texas at Austin) and Dr. Daniel Sturmer (Ph.D., University of Nevada, Reno, 2012; currently at Shell Exploration and Production Company), respectively. The late Dr. John D. Cooper, CSUF Emeritus Professor, guided the initial description and interpretation of the sedimentary protoliths. It is the authors’ great sorrow that “Coop” could not participate in the final synthesis. CSUF advanced petrography students and alumnae Crystal Castellanos, Michelle Gevedon, Kimberly Nepsa, and Michelle Slopko Kane are thanked for their contributions. Field support from Tish Butcher, Jeff Knott, and Steve Turner is gratefully acknowledged. Use of the Pomona College XRF laboratory, which was partially funded by NSF grant DUE-CCLI 0942447, was made possible through the hospitality of Jade Star Lackey. This project would not have been possible without the cooperation and support of the eight landowners of the Goldstein Peak Formation who made the senior author and her students feel welcome. In a class by herself is rancher Barbara Chrisman, who has supported geologic fieldwork and education in the shadow of Goldstein Peak for over 20 years. Sincere thanks, and a debt of gratitude, are owed to them all.