Subduction Initiation Along Transform Faults: the Proto-Franciscan Subduction Zone

Subduction initiation along transform faults: The proto-Franciscan subduction zone

John W. Shervais1* and Sung Hi Choi2 1DEPARTMENT OF GEOLOGY, UTAH STATE UNIVERSITY, 4505 OLD MAIN HALL, LOGAN, UTAH 84322-4505, USA 2DEPARTMENT OF GEOLOGY AND EARTH ENVIRONMENTAL SCIENCES, CHUNGNAM NATIONAL UNIVERSITY, DAEJEON 305-764, SOUTH KOREA

ABSTRACT

The initiation of subduction is a process that cannot be observed directly but must be inferred from the rock record after subduction is well established. There are many approaches possible to infer the origin of subduction zones that are still active, but paleosubduction zones present special challenges, since their geodynamic setting can no longer be directly observed. In this study, we examine evidence for subduction initiation of the proto-Franciscan subduction zone along a transform fault, based on a subduction initiation origin for the Coast Range ophiolite, and on the Tehama-Colusa serpentinite mélange, which underlies the ophiolite and separates it from high-pressure/temperature metamorphic rocks of the Franciscan complex. The Coast Range ophiolite consists of volcanic, plutonic, and mantle components, each of which contains elements that refl ect subduction initiation or hydrous melting within a subduction-zone setting. The volcanic assemblage includes forearc basalts and boninites, as well as more evolved calc-alkaline rocks; the plutonic complex contains intrusive suites that refl ect this same range of parent magmas. Peridotites of the mantle section include both abyssal-like and refractory peridotites formed by hydrous decompression melting. The Tehama-Colusa serpentinite mélange consists of blocks of basalt, chert, sedimentary rocks, and peridotite (harzburgite and lherzolite) in a sheared serpentinite matrix. The mélange matrix represents hydrated refractory peridotites with forearc affi nities, and blocks within the mélange consist largely of upper-plate lithologies (harzburgite, arc volcanics, and arc-derived sediments). Lower-plate blocks within the mélange include oceanic basalts and chert with rare blueschist and amphibolite. The abyssal peridotites have low pyroxene equilibration temperatures that are consistent with formation in a fracture-zone setting. However, the current mélange refl ects largely upper-plate lithologies in both its matrix and its constituent blocks. We propose that the proto-Franciscan subduction zone nucleated on a large offset transform fault or fracture zone that evolved into a subduction-zone mélange complex. The nucleation of subduction zones along former transform boundaries has long been proposed for both modern arc systems and for the Franciscan–Coast Range ophiolite system. Our data support this interpretation and document more fully how this mechanism is expressed by mixing within the evolving serpentinite mélange.

LITHOSPHERE; v. 4; no. 6; p. 484–496 | Published online 14 December 2011 doi: 10.1130/L153.1

INTRODUCTION stable oceanic lithosphere, or adjacent to buoyant continental crust (e.g., Casey and Dewey, 1984; Leitch, 1984; Stern and Bloomer, 1992; Stern, The process of subduction initiation cannot be observed directly but 2004). In either case, the ultimate result is a subduction zone that underlies must be inferred from the rock record after a subduction zone is estab- a volcanic arc, where the arc rests on older oceanic crust or continental lished. “Induced” subduction initiation (Hall et al., 2003; Stern, 2004) crust (Tatsumi and Eggins, 1995). Investigations of early arc volcanism results from far-fi eld plate stresses that force convergence across a zone over the past two decades have shown that primitive early arc volcanics of weakness within a plate during plate-boundary reorganizations (e.g., comprise the same rock assemblages as most ophiolites, and they are com- the Maquarie-Puysegur-Fiordland system near New Zealand; Ruff et al., monly preserved in highly extended forearc regions, which are interpreted 1989) or in response to a collisional event between an existing subduction to represent the products of spontaneous subduction initiation (Bloomer et zone and continental or unusually thick oceanic crust (e.g., Ontong Java al., 1995; Stern and Bloomer, 1992; Hawkins, 2003; Reagan et al., 2010). Plateau collision with the Solomon Islands Trench; Cooper and Taylor, Most extensive ophiolite terranes are believed, on the basis of their 1985). In contrast, “spontaneous” subduction initiation occurs most com- lavas and mantle residues, to have formed in, or passed through, a supra- monly when stable ocean lithosphere, which forms by anhydrous decom- sub duction-zone environment (e.g., Miyashiro, 1973; Alabaster et al., pression melting at mid-oceanic spreading centers and essentially “fl oats” 1982; Shervais, 1982; Pearce et al., 1984; Shervais and Kimbrough, 1985; on top of the underlying asthenosphere, becomes gravitationally unstable Metcalf and Shervais, 2008). While the specifi c environment (backarc, and begins to sink back into the mantle, either adjacent to gravitationally arc, and forearc) is often debated, it is clear that the volcanic rock series found in ophiolites most closely corresponds to lavas now found within the forearc region of active arc terranes (e.g., Metcalf and Shervais, 2008). Further, it has been proposed that at their time of formation, there is little *E-mail: [email protected]. or no evidence for the existence of a volcanic arc (Stern and Bloomer, Editor’s note: This article is part of a special issue titled “Initiation and Termina- tion of Subduction: Rock Record, Geodynamic Models, Modern Plate Boundaries,” 1992; Pearce, 2003; Metcalf and Shervais, 2008). edited by John Shervais and John Wakabayashi. The full issue can be found at Shervais (2001) showed that suprasubduction-zone ophiolites display http://lithosphere.gsapubs.org/content/4/6.toc. a consistent sequence of events during their formation and evolution, sug-

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gesting that they form in response to processes that are common to all such 2004; Metcalf and Shervais, 2008; Whattam and Stern, 2011), the specifi c ophiolites. This sequence includes: (1) birth—formation of the ophiolite setting or circumstances of this event are commonly unclear. Our second above a nascent or reconfi gured subduction zone; (2) youth—continued focus here is on the evidence provided by serpentinite mélange zones, melting of refractory asthenosphere (depleted during birth) in response to which are commonly associated with suprasubduction-zone ophiolites. fl uid fl ux from the subducting slab; (3) maturity—onset of semistable arc The petrologic and geochemical character of blocks within these mélange volcanism; (4) death—the sudden demise of active spreading and ophio- zones provides clues regarding their origin, which may be unrelated to the lite-related volcanism; and (5) resurrection—emplacement by obduction adjacent ophiolite, or may refl ect a linkage among the mélange, subduc- onto a passive margin or accretionary uplift with continued subduction. tion initiation, and ophiolite formation. This sequence of events is similar to that inferred by Reagan et al. (2010) We take as our case example the Coast Range ophiolite of Califor- from their detailed studies of the Izu-Bonin-Mariana forearc, with tholei- nia, which has been studied extensively for over three decades (Hopson itic “forearc basalts” preceding boninites, and with both forming during et al., 1981, 2008; Shervais and Kimbrough, 1985; Shervais, 1990, 2001; an episode of rapid extension in the forearc prior to establishment of the Coleman, 2000; Shervais et al., 2005a, 2005b). The Coast Range ophiolite calc-alkaline volcanic arc. is closely linked to the Franciscan assemblage—one of the world’s most Our primary focus here is on the fi rst two stages proposed by Shervais intensely studied subduction complexes and a model for convergent bound- (2001): (1) birth, characterized by forearc basalts (Reagan et al., 2010), ary plate processes (Bailey et al., 1964; Blake and Jones, 1974; Blake et represented by a variety of early arc tholeiites with strong mid-ocean- al., 1982; Wakabayashi, 1999; Ernst, 1993). In northern California, the ridge basalt (MORB)–like characteristics, and (2) youth, characterized by ophiolite is separated from younger rocks of the Franciscan assemblage by high-Mg basalts and andesites of the boninite suite, which are among the the Tehama-Colusa mélange—a serpentinite-matrix mélange containing most depleted volcanic rocks on Earth (e.g., Metcalf and Shervais, 2008). blocks of peridotite, metabasalt, chert, and high-grade metamorphic rocks These stages are thought to represent the onset and development of sub- that has been interpreted as an oceanic fracture-zone assemblage (Hopson duction initiation, followed by the transition to stable subduction (stage and Pessagno, 2005) linked to subduction initiation (Choi et al., 2008b). three: maturity and calc-alkaline volcanism). Ophiolites that represent the fi rst two stages in this progression (birth-youth) are interpreted to have GEOLOGIC SETTING formed during a subduction initiation event, regardless of their current apparent setting (Fig. 1). Further, these stages can be established by care- The Coast Ranges of California comprise a complex orogen thought ful mapping within the plutonic series of the ophiolites (Shervais et al., to have formed in the forearc region of the Sierra Nevada arc (or farther 2004), and they can be inferred from highly refractory peridotite assem- south, e.g., Wright and Wyld, 2007) during the Mesozoic and Paleogene blages within the underlying mantle tectonites (e.g., Choi et al., 2008a, (Fig. 2; Shervais et al., 2004; Hopson et al., 2008). This orogen is dis- 2008b; Jean et al., 2010). tinct from the classic collisional orogens (e.g., the Appalachians, Alps, or Although suprasubduction-zone ophiolites are commonly linked to Himalayas) because the events preserved here occurred within an active subduction initiation (Stern and Bloomer, 1992; Shervais, 2001; Stern, convergent system that did not experience a major continental collision—

Spreading Trench: Thin crust formed by rapid extension in response to slab sinking center hinge rollback Depth Arc tholeiites (FAB) —> boninites —> calc-alkaline volcanics (km)

20 Extending lithosphere 40 Mantle wedge melt zone 12-30 kb (40-100 km) 60 Figure 1. Schematic diagram of a nascent “spontaneous” DehydrationAqueous reactions fluids subduction zone; no vertical exaggeration. Sinking oceanic lithosphere displaces hot asthenosphere, which fl ows up 80 Inflow of hot into the extensional gap created by the sinking lithosphere. asthenosphere The asthenosphere undergoes decompression melting with 100 minor involvement of slab-derived fl uids to form mid-ocean- ridge basalt (MORB)–like arc tholeiitic basalts. As slab sink- 120 ing proceeds, more fl uids are released and fl uid-enhanced melting of the previously depleted asthenosphere occurs, forming boninite suite magmas. Eventually stable subduc- 140 tion ensues, leading to normal calc-alkaline magmatism. Slab sinking gravitationally 160

0 km 100 km 200 km 300 km

Distance from trench (km)

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Eureka all of the deformation and metamorphism found here formed entirely 124° Accreted 122° within a subduction-zone setting. This orogen documents the complexity terranes that may arise within any active margin prior to a continent-continent collision (or arc-continent collision), and it shows that not all of the structures and metamorphic features found in collisional orogens are necessarily related to that collision (e.g., Yarlung-Tsangpo suture; Ratschbacher et al., 40°N NE 40°N 1992; Guilmette et al., 2008). EC The Coast Range orogen of California is composed of three main

SW elements: the Upper Jurassic to Upper Cretaceous Great Valley Group, deposited within a complex forearc basin; the Middle Jurassic Coast SFV Range ophiolite, which forms the basement upon which the Great Valley Group was deposited; and the Franciscan assemblage, which represents

SAF Arc complex the accretionary prism of the Cordilleran subduction zone (Fig. 3). Here, Forearc basin we discuss two of those elements—the Coast Range ophiolite and the Franciscan assemblage—and the Tehama-Colusa mélange, which sepa- 38°N 38°N rates these units in the northern Coast Ranges. Post-Mesozoic SF Coast Range Ophiolite Franciscan accretionary DP complex Eastern belt The Coast Range ophiolite is a linear belt of felsic, mafi c, and ultra- Central belt mafi c rocks exposed primarily along the eastern fl ank of the California Coast Ranges, but also as scattered locales west of the Salinia block in the Coastal belt Salinia Ll SNF western Transverse Ranges and farther north along the coast (Fig. 2). It is 36°N one of the most extensive ophiolite terranes in North America and has long 36°N been central to our understanding of Cordilleran tectonics. Nonetheless, its Modoc plateau origin is controversial and three primary hypotheses have been advanced: Salinia SAF (1) formation at a mid-ocean-ridge spreading center at low paleolatitudes, Great Valley Group CR and subsequent rapid drift northward to collide with North America (Hop- Coast Range ophiolite son et al., 1981, 2008; Pessagno et al., 2000); (2) formation as a backarc Sierra Nevada PS basin behind an east-facing volcanic arc that collided with North America Klamath terranes SB during the Late Jurassic Nevadan orogeny (Godfrey and Klemperer, 1998; 124°W 122°W 120°W Ingersoll, 2000); and (3) formation by forearc or intra-arc rifting along the western margin of North America, in response to nascent or renewed Figure 2. Geologic map of California showing main elements of the Coast subduction of oceanic plates beneath North America (Shervais and Kim- Range geology. Modifi ed from Hopson et al (1981). SF—San Francisco, brough, 1985; Shervais, 1990, 2001; Stern and Bloomer, 1992). The con- SB—Santa Barbara, SAF—San Andreas fault, SNF—Sur-Nacimiento fault. Major Coast Range ophiolite localities (N-S): EC—Elder Creek, SFV— sensus of most recent studies support the suprasubduction-zone model Stonyford (Black Diamond Ridge), DP—Del Puerto Canyon, Ll—Llanada, (Stern and Bloomer, 1992; Shervais et al., 2004, 2005a; Shervais, 2008; CR—Cuesta Ridge, PS—Point Sal.

Central belt melange Great Valley Group Coastal belt Prehnite-pumpellyite facies with Eastern belt forearc basin zeolite facies low-grade (subgreenschist, greenschist facies) Blueschist facies underlain by & high-grade knockers Coast Range ophiolite Trench slope basins coherent slabs Subducting and melange oceanic slab Figure 3. Cross section of the Mesozoic active margin of California. Structural accretion and metamorphic ages become younger to the Coast Range ophiolite & west; metamorphic grade Coast Range Tehama-Colusa mélange becomes generally lower fault from west to east. Modiﬁ ed Central belt Great 10 km from Shervais (2006), after mélange Valley Group Blake et al (1985). 10 km Eastern belt Coastal belt Pickett Peak terrane

Eastern belt Oceanic slab Yolla Bolly terrane

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Choi et al., 2008a; Jean et al., 2010), but models based on backarc basins volcanic arc tholeiite basalts (similar to MORB in trace-element com- and oceanic spreading centers still persist (e.g., Hopson et al., 2008). position), overlain by boninites, or boninitic basalts (which also form The Coast Range ophiolite is composed of three major components: dikes that crosscut the earlier tholeiites). The MORB-like arc tholeiites the volcanic-hypabyssal upper crust and its overlying section of volcano- are similar to the forearc basalts of Reagan et al. (2010). Calc-alkaline pelagic sediments, the plutonic lower crust, and the underlying refractory volcanics, which occur at many locales (e.g., Del Puerto, Llanada, Elder peridotites, which represent the upper mantle upon which the crust was Creek), form a fractionation series of basalt-andesite-dacite-rhyolite built after the extraction of partial melts parental to the crust (Hopson et (Fig. 4). At Del Puerto, these include a younger series of hornblende- al., 1981, 2008; Shervais et al., 2004). The crustal components display bearing volcaniclastics called the Lotta Creek tuff (Evarts et al., 1999), the developmental cycle described by Shervais (2001) that has been while at Elder Creek, the calc-alkaline volcanics are found only in a sed- interpreted to refl ect a consistent sequence related to subduction initia- imentary breccia called the Crowfoot Point Breccia, which overlies the tion and the subsequent transition toward more normal arc activity. The ophiolite disconformably (Hopson et al., 1981, 2008; Robertson, 1990; mantle rocks preserve evidence for the early part of this cycle, as well as Shervais et al., 2004). Relationships at Llanada are similar to Del Puerto, for events that predate ophiolite formation (Choi et al., 2008b). with extrusive volcaniclastics and intrusive keratophyre sills (Giaramita et al., 1998; Hopson et al., 2008). Volcanic-Hypabyssal Upper Crust and Volcano-Pelagic Sediments The upper crust at most Coast Range ophiolite localities consists of Plutonic Lower Crust mafi c volcanic rocks, which commonly transition upward into more- The lower crust of the Coast Range ophiolite can be divided into three evolved intermediate-felsic volcanics, and an overlying cover of ash- suites: (1) a mafi c cumulate series of layered gabbros with associated rich radiolarian cherts and tuffs. The volcanic rocks are dominated by dunites and isotropic gabbros, (2) an intrusive series of ultramafi c-mafi c

20.00

A FAB-IBM Stage 1 Stage 2 15.00 Stage 3 FAB Stage 4

10.00 FeO* (wt%)

5.00

0.00 Figure 4. Variation diagrams for ophiolite vol-

45.00 50.00 55.00 60.00 65.00 70.00 75.00 80.00 canics compared to Mariana forearc: (A) SiO2 vs FeO* (total Fe as FeO) and (B) MgO vs Ti SiO (wt%) O . Mariana forearc basalt (FAB) data from 2 2 Reagan et al (2010). CRO data from Shervais 3.0 (2001, 2008). IBM—Izu-Bonin-Mariana arc.

B FAB-IBM 2.5 Stage 1 Stage 2 Stage 3 2.0 Stage 4

1.5 FAB (wt%) 2

TiO 1.0 boninitic 0.5 boninites 0.0 0.0 5.0 10.0 15.0 20.0 MgO (wt%)

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cumulates made up of wehrlite, olivine clinopyroxenite, and isotropic gab- et al., 1972). The early to mid-Cretaceous Eastern belt is subdivided into bro (±dunite), and (3) a later intrusive series of tonalite-trondhjemite-dio- the Pickett Peak terrane (South Fork Mountain Schist, Valentine Spring rite (TTD) that postdates the fi rst two plutonic suites. Cumulate gabbros Formation; Worrall, 1981) and the subjacent Yolla Bolly terrane (Blake of suite 1 are dominant and typically form over 60% of the lower-crustal et al., 1982). The Pickett Peak terrane consists of blueschist-facies mud- section. These rocks are typical of ophiolite layered gabbro sequences, stone and basalt (South Fork Mountain Schist) and mudstone-graywacke with ~60%–70% modal plagioclase and 30%–40% modal pyroxene or (Valentine Spring Formation), forming coherent sheets up to hundreds of pyroxene and amphibole; olivine is relatively rare. Augite is the dominant meters thick transposed along low-angle thrust faults (Blake et al., 1982). pyroxene, although hypersthene may occur as well in place of hornblende. The Yolla Bolly terrane consists of blueschist-facies mudstone and gray- Suite 2 wehrlites, clinopyroxenites, and isotropic gabbros are intrusive wacke, with minor basalt and chert, in coherent thrust sheets, including into the layered gabbros of suite 1, documented by massive wehrlite- metamorphosed mélange (Fig. 3; Bailey et al., 1964; Blake and Jones, pyroxenite dikes that crosscut layering in the gabbro, and by xenoliths 1974; Wakabayashi, 1992, 1999; Ernst, 1993). The Coastal belt, which of layered gabbro in wehrlite and suite 2 isotropic gabbro (Hopson et al., ranges in age from mid-Cretaceous through Miocene, consists of west- 1981; Shervais, 2001; Shervais et al., 2004). The TTD suite (suite 3) forms verging thrust slices that get younger to the west, comprising both coher- small intrusive bodies, thick (10–15 m) dikes, and fl at-lying, kilometer- ent graywacke and broken formations of arkose, mudstone, and conglom- scale plutons that intrude all of the older plutonic suites and parts of the erate (Fig. 3; McLaughlin et al., 1982; Blake et al., 1985). The Central overlying volcanic-hypabyssal complex (Shervais, 2001, 2008; Shervais belt mélange contains blocks of graywacke, greenstone, serpentinite, et al., 2004). Radiometric U-Pb zircon ages for the Coast Range ophiolite chert, limestone, blueschist, eclogite, and garnet amphibolite, in a matrix represent this fi nal intrusive suite; 238U-206Pb ages typically range from of fi nely comminuted micrograywacke siltstone/shale (Bailey et al., 1964; ca. 172 Ma to ca. 162 Ma (Shervais et al., 2005a; Hopson et al., 2008; Mat- Blake and Jones, 1974; Blake and Wentworth, 1999). Large tracts of tinson and Hopson, 2008), although some locales contain a younger suite arkosic wackes interpreted as slope basin deposits form mappable units of hornblende-bearing quartz diorites and keratophyres with 238U-206Pb within the mélange (Becker and Cloos, 1985). and 40Ar/39Ar ages of ca. 150 Ma (Evarts et al., 1999; Hopson et al., 2008). Tehama-Colusa Mélange Subcrustal Mantle Peridotites The crustal series described here is underlain structurally by mantle The Tehama-Colusa serpentinite mélange is restricted to the northern peridotites—including lherzolite, harzburgite, and dunite—that refl ect Coast Ranges, where it separates the Middle Jurassic Coast Range ophio- increasing melt extraction and progressively more refractory composi- lite from Lower Cretaceous rocks of the Franciscan assemblage Eastern tions (Choi et al., 2008a; Jean et al., 2010). Contacts with the overlying belt (Huot and Maury, 2002; Hopson and Pessagno, 2005; Shervais et al., crustal section are rarely exposed and, where seen, are generally faulted. 2011). The Tehama-Colusa mélange extends for over 160 km from Wilbur In the northern Coast Ranges, these blocks are typically associated with Springs to the North Fork of Elder Creek (Fig. 2). The mélange consists of the underlying Tehama-Colusa serpentinite mélange, whereas in the Dia- a sheared serpentinite matrix with tectonic blocks of abyssal and refractory blo Range and west of the Sur-Nacimiento fault, the peridotites directly peridotite, lower-crustal plutonic rocks (cumulate gabbro, wehrlite, and underlie the crustal sequence. Harzburgites, with lesser dunite, are the diorite), low-grade metavolcanic rocks, high-grade metamorphic blocks most common lithologies at all locations where peridotites crop out; how- of blueschist, amphibolite, and garnet amphibolite, volcaniclastic sand- ever, lherzolite is found locally at Cuesta Ridge and forms a large, kilome- stones of uncertain provenance, and foliated metasediments (Jayko and ter-scale block near Stonyford (Fig. 2). Jean et al. (2010) have shown that Blake, 1986; Hopson and Pessagno, 2005; Shervais et al., 2011). Many the harzburgites represent up to 23% melt extraction under decompressing of these blocks were derived from the overlying Coast Range ophiolite, pressure conditions, fi rst in the garnet fi eld and then in the spinel facies. e.g., the lower-crustal gabbros, wehrlites, and diorites (which have U-Pb The high fraction of melting and the modal compositions are consistent zircon ages identical to diorites of the Coast Range ophiolite; Shervais et with hydrous melting in the mantle wedge of a subduction zone, a con- al., 2005a), and many of the low-grade metavolcanic blocks, which are clusion supported by the high fl uid mobile element concentrations and identical in composition and grade to volcanic rocks in the ophiolite. It evolved isotopic compositions of clean clinopyroxene mineral separates is inferred that the volcaniclastic sandstones were derived from the Coast (Choi et al., 2008a, 2008b). In contrast, the lherzolites have mineral com- Range ophiolite as well, and that they may be correlated with the Crowfoot positions similar to abyssal peridotites formed at mid-ocean-ridge spread- Point Breccia (Seymore-Simpson, 1999), whereas the foliated metasedi- ing centers, but they are still characterized by high fl uid mobile element ments have been correlated with the Galice Formation (Jayko and Blake, concentrations and evolved isotopic compositions (Choi et al., 2008a, 1986). These rocks document a close association between the mélange 2008b; Jean et al., 2010). and the overlying ophiolite, which requires that the mélange and ophiolite formed within the same tectonic realm and that mélange formation contin- Franciscan Assemblage ued after formation of the ophiolite crustal assemblages. Peridotite blocks of the Tehama-Colusa mélange can be classifi ed into The Franciscan assemblage formed during prolonged subduction of two groups, as discussed previously for the entire Coast Range ophiol- oceanic lithosphere beneath the western margin of North America in the ite. The dominant group is highly refractory (harzburgite and dunite) and late Mesozoic and early Tertiary (Bailey et al., 1964; Hsü, 1968; Ernst, refl ects extensive melt extraction in a hydrous melting environment (Huot 1993; Wakabayashi, 1999). It extends from Eureka to the southern Diablo and Maury, 2002; Choi et al., 2008a, 2008b; Jean et al., 2010). The Black Range east of the San Andreas fault system, and from Big Sur to the west- Diamond Ridge massif near Stonyford consists of abyssal-like lherzolite ern Transverse Ranges west of the Sur-Nacimiento fault, encompassing an that has shown limited melt extraction (Choi et al., 2008a, 2008b; Jean et area ~700 km long and up to 200 km wide (Fig. 2). The Franciscan assem- al., 2010). The refractory harzburgites have compositions consistent with blage includes three tectonic belts that generally become progressively extraction of partial melts to form the overlying ophiolite massif; the prove- younger (structural and metamorphic ages) from east to west: the Eastern nance of the fertile lherzolites is uncertain, but their trace-element and isoto- belt, the Central belt, and the Coastal belt (Bailey et al., 1964; Berkland pic compositions suggest residence within a subduction-zone environment

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for at least part of its existence. Finally, the mélange matrix has been shown positional range (basalt, basaltic andesite, and andesite; Fig. 4) that have to contain relict spinel grains with a wide range of Cr/Al ratios, consistent major- and trace-element characteristics transitional to MORB (Fig. 5). with a mixed provenance of highly refractory harzburgite-dunite and more These volcanic rocks are equivalent to the so-called “forearc basalts” of fertile lherzolite (Huot and Maury, 2002; Shervais et al., 2011). the Mariana forearc (Reagan et al., 2010), and they occupy a similar setting—they predate all other volcanic arc basalts and underlie the later METHODS boninitic lavas (e.g., Shervais, 2001). The forearc basalts are followed by boninites (extremely depleted high-Mg andesites) and boninitic lavas Volcanic rock compositions were compiled from published sources (somewhat less-depleted or more-evolved volcanics directly related to and plotted on standard petrogenetic diagrams to allow comparison with boninite magmatism). The boninites and boninitic lavas correspond to other settings. We also present new secondary ion mass spectrometry stage two (youth) of Shervais (2001). These magmas form in response to (SIMS) analyses of in situ clinopyroxene grains from gabbros, wehrlites, extremely high melt fractions—a direct result of continued decompression and pyroxenites of the Elder Creek ophiolite, along with calculated equi- melting in the presence of high fl uid phase fl uxes. They are characterized librium magma compositions based on pyroxene-melt partition coeffi - by higher MgO and Cr contents, and by lower concentrations of Ti, Y, and cients. The SIMS analyses were carried out on the Cameca 4f ion micro- other moderately incompatible elements (Figs. 4 and 5). probe at the University of New Mexico. Finally, we compiled published Stage three lavas are calc-alkaline rocks thought to be related to the data for mantle peridotites that underlie the ophiolite, or that are part of the onset of stable subduction (maturity), but the more magnesian members of Tehama-Colusa mélange, and separate it from the Franciscan assemblage, this suite may form in response to fractionation of primitive boninitic mag- in order to deduce the original setting that led to subduction initiation. mas. Finally, in the Coast Range ophiolite, stage four lavas are MORBs formed in response to propagation of a backarc spreading center into an RESULTS arc (e.g., Josephine ophiolite; Harper, 2003) or collision of the subduction zone with an actively spreading oceanic ridge (death; Shervais, 2001). Volcanic Rock Compositions Plutonic Rock Parent Magmas Shervais (2001) divided volcanic rocks of the Coast Range ophiolite into four main stages, based on composition, relative age, and occurrence. Detailed mapping and fi eld work carried out over the past 20 yr have The fi rst stage (birth) includes volcanic arc tholeiites with limited com- shown that many ophiolite plutonic suites—once thought to represent

500 VAB 20 450 FAB-IBM FAB melting curve Stage 1 Stage 2 400 1000 60% Stage 3 40% Stage 4 MORB 20% 10% 5% 350 50 FAB 300 MORB 100 250 V (ppm)

Alkali Cr (ppm) 200 basalt

150 FAB-IBM Stage 1 10 100 Stage 2 Stage 3 Stage 4 IAB 50 AB 0 1 0 5000 10,000 15,000 20,000 1020 30 40 50 60 80 100

Ti (ppm) Y (ppm) Figure 5. Trace element discrimination diagrams for CRO (Coast Range ophiolite) and Mariana forearc volcanics: (A) Ti vs V (after Shervais 1982); the constant ratio lines 20 and 50 generally separate arc rocks (<20) from MORB (mid-ocean-ridge basalt) (20–50) and alkali basalts (>50). (B) Y vs Cr (after Pearce et al 1984); curve in upper part of diagram is melting curve for MORB-source mantle, marked in percent melt increments. In general, MORB refl ects lower percent melts than arc basalts due to absence of fl uid fl ux. Data sources same as Figure 4. IBM—Izu-Bonin-Mariana arc; FAB—forearc basalt; VAB—volcanic arc basalt; IAB—island arc basalt.

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TABLE 1. MAJOR- AND RARE EARTH ELEMENT ANALYSES OF CALCIC PYROXENE FROM PLUTONIC ROCKS OF THE ELDER CREEK OPHIOLITE, CALIFORNIA EC-17-2 EC-21-6 EC-20-2 EC105-3 EC131-1 EC141-1 EC104-1 EC131-2 EC144-1 EC135-1 EC91-1 EC45-1 EC98-1 Relative error Lithology W W P FP CG CG CG CG CG CG CG IG IG 1σ

SiO2 53.25 52.42 52.60 51.31 51.79 51.15 51.10 50.86 51.30 51.70 50.01 49.58 50.32

TiO2 0.09 0.12 0.20 0.23 0.22 0.34 0.40 0.41 0.27 0.40 0.40 0.29 0.28

Al2O3 1.24 2.24 2.40 2.88 1.84 2.77 2.80 2.62 2.74 1.80 3.96 3.13 2.92 FeO 2.14 2.99 3.30 5.32 6.46 6.57 6.80 7.20 7.30 7.40 7.65 13.28 11.83 MnO 0.07 0.10 0.10 0.14 0.21 0.18 0.20 0.22 0.18 0.10 0.16 0.34 0.32 MgO 17.62 17.24 17.10 16.53 15.53 15.57 15.40 15.47 15.47 15.10 14.93 14.67 14.54 CaO 24.81 23.54 22.70 22.37 22.60 22.58 22.10 21.64 22.08 22.30 21.70 17.67 19.09

Na2O 0.06 0.15 0.20 0.46 0.21 0.24 0.20 0.28 0.28 0.20 0.26 0.20 0.21

K2O 0.002 0.000 0.000 0.005 0.001 0.000 0.000 0.000 0.001 0.000 0.002 0.002 0.000

Cr2O3 0.336 0.607 0.700 0.123 0.036 0.160 0.100 0.100 0.022 0.000 0.002 0.011 0.011 Total 99.61 99.41 99.20 99.37 98.91 99.55 99.10 98.80 99.66 99.20 99.07 99.18 99.51 mg# 93.62 91.13 90.23 84.71 81.08 80.86 80.15 79.30 79.07 78.44 77.68 66.32 68.66

La 0.03091 0.11963 0.14364 0.17013 0.32246 0.26500 0.11948 0.28287 0.26814 0.15385 0.15835 0.61467 0.26817 8.6% Ce 0.16934 0.50701 0.60520 0.86675 2.05152 1.35176 0.70245 1.83263 1.58766 0.92533 0.85896 1.92281 0.81234 4.4% Nd 0.24800 0.65703 0.84118 1.58561 3.70958 2.45766 1.38240 3.61893 2.87213 1.47747 1.79000 2.77182 1.25514 5.2% Sm 0.13349 0.32399 0.39944 1.00081 2.13968 1.57527 0.79889 2.08570 1.62655 0.82187 1.28620 1.64766 0.86582 5.8% Eu 0.05647 0.13982 0.16269 0.38109 0.77387 0.56290 0.40116 0.76862 0.63438 0.30095 0.50963 0.62444 0.38481 6.3% Gd* 0.16456 0.41296 0.52356 1.21876 2.73188 1.97571 1.12652 2.71157 2.06142 1.15890 1.67215 2.27992 1.08146 Dy 0.18519 0.48298 0.63112 1.35131 3.20110 2.26883 1.45851 3.24052 2.39558 1.50042 1.99832 2.90068 1.23602 4.3% Er 0.12892 0.27429 0.32966 0.83176 1.86116 1.30937 0.80300 1.90813 1.35993 0.79004 1.16463 1.95959 0.76336 4.5% Yb 0.10295 0.24250 0.31843 0.67769 1.59820 1.15077 0.78712 1.66372 1.20105 0.76034 0.92405 1.86970 0.63742 3.4% Ce/Yb 1.64 2.09 1.90 1.28 1.28 1.17 0.89 1.10 1.32 1.22 0.93 1.03 1.27 Note: Major elements in weight percent oxide, rare earth elements in ppm. Gd* is interpolated from chondrite normalized Sm and Dy concentrations. Most represent average of 2–6 spot analyses. W—wehrlite, P—pyroxenite, FP—feldspathic pyroxenite, CG—cumulate gabbro, IG—isotropic gabbro.

simple cumulate piles—are in fact polymagmatic complexes composed of trondhjemite-quartz diorite, which may grade into isotropic gabbro as at least two or three distinct magmatic suites (Juteau et al., 1988; Hebert modal quartz declines (Shervais, 2001, 2008; Shervais et al., 2004). and Laurent, 1990; Laurent, 1992). The most common suites are similar to Crosscutting relationships within the plutonic complex clearly estab- those found in the Coast Range ophiolite: (1) a primary suite of cumulate lish the order of intrusion, which is consistent from one ophiolite to dunite and layered gabbros composed of plagioclase and clinopyroxene, another, and this sequence of intrusion is similar to the eruptive order with less common olivine, orthopyroxene, or hornblende; (2) a later intru- noted in the volcanic rocks. This implies that the early cumulate gab- sive suite of wehrlites, pyroxenites, and primitive isotropic gabbros; and bros are related to the forearc basalts, that the wehrlite-pyroxenite suite (3) a fi nal suite of calc-alkaline dikes and plutons composed of tonalite- is derived from the boninitic suite magma, and that the tonalite-trondhjemite-quartz diorite suite is derived from the same magmas as the calc-alkaline volcanic rocks. We can assess these potential relationships by calculating the rare earth element (REE) concentration of magmas 1000 in equilibrium with cumulate pyroxene in the gabbros and wehrlites, Isotropic gabbro using pyroxene SIMS analyses (Table 1). In Figure 6, we compare these EQ melt cumulate gabbro equilibrium melts with the forearc basalts and boninites of Reagan et EQ melt wehrlite-clinopyroxenite 100 al. (2010), and with two isotropic gabbros related to the stage-one layered gabbros. As can be seen in the fi gure, the isotropic gabbros and the FAB equilibrium melts with pyroxenes from the layered gabbros display REE concentrations similar to forearc basalts. In contrast, pyroxenes from 10 the wehrlites and clinopyroxenites are in equilibrium with melts that have REE concentrations similar to boninitic melts—the overall REE concentrations are lower, but they are enriched in light (L) REE relative 1 to heavy (H) REEs, probably as a result of slab-derived fl uid metasoma- Boninite suite tism (Fig. 6). Subcrustal Mantle Peridotites and the Tehama-Colusa Mélange 0.1 La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu The mixing of volcanic, plutonic, and sedimentary rocks of Coast Figure 6. Chondrite-normalized rare earth element diagram with con- Range ophiolite (upper plate) provenance with volcanic and sedimentary centrations for melts calculated to be in equilibrium (EQ) with pyroxene from cumulate gabbro and wehrlite of Coast Range ophiolite. Also shown rocks of oceanic (lower plate) provenance has been well documented in are two isotropic gabbros (melt compositions) and volcanic rocks of the the Tehama-Colusa mélange (Shervais et al., 2011). We focus here on Mariana forearc (FAB—forearc basalts and boninites; Reagan et al, 2010). peridotite blocks within the serpentinite matrix mélange and on massive

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peridotites that directly underlie the ophiolite—which also exhibit mixed 100 provenance behavior. 90 Peridotites associated with the Coast Range ophiolite defi ne two dis- Abyssal spinels tinct groups. The dominant group found at all locations includes highly 80 Forearc spinels refractory harzburgite and dunite, often cut by veins or dikes of orthopy- 70 roxenite (Choi et al., 2008a). Pyroxenes in the refractory harzburgites have spoon-shaped REE patterns (strongly depleted middle REEs with slightly 60 enriched LREEs) and high concentrations of fl uid mobile incompatible 50 trace elements (Jean et al., 2010). Spinels in these harzburgites, dunites, Mg# 40 and orthopyroxenites have high Cr#’s (100 × Cr/[Cr + Al]) that corre- CRO lherzolite spond to the compositional range of forearc peridotites (Fig. 7). The less 30 CRO harzburgite CRO dunite common group found at a few locations is made up of relatively fertile 20 CRO pyroxenite lherzolite and clinopyroxene-harzburgite with aluminous spinels (Fig. 7; Mélange matrix Choi et al., 2008A; Jean et al., 2010). Spinels recovered from the scaly 10 serpentinite matrix of the mélange (Huot and Maury, 2002) are mostly 0 refractory Cr-rich spinels, with a few more aluminous spinels, showing 0102030405060708090100 that the matrix is formed largely by the shearing of upper plate (forearc) Cr# peridotite (e.g., Shervais et al., 2011). Figure 7. Quadrilateral plot of spinel compositions (Mg# vs Cr#) from Figure 8 shows the Cr# of clinopyroxene as a function of equilibrium ophiolite peridotites compared to spinels from abyssal peridotites and two-pyroxene temperatures of the Coast Range ophiolite peridotites cal- refractory fore-arc peridotites. Data from Shervais et al 2011. CRO—Coast culated using the geothermometer of Brey and Köhler (1990). Informa- Range ophiolite. tion also shown for comparison includes the available data for abyssal peridotites associated with major fracture zones (Vema, Owen, Roman- che) and normal, small-offset transforms. Major fracture-zone peridotites are characterized by lower equilibration temperatures at a given Cr# than normal, small-offset transform peridotites (Fig. 8). All Coast Range ophiolite peridotite pyroxenes have low equilibration temperatures. The refractory Coast Range ophiolite peridotites fall outside the fi eld for all abyssal peridotites, whereas the fertile Coast Range ophiolite peridotites 600 (lherzolites) have temperatures that correspond to large-offset transform faults/fracture zones (Choi et al., 2008b). The occurrence of low equilibration temperatures in mantle associated with large-offset fracture zones 800 VFZ may refl ect hydrothermal cooling of the crust and mantle to deeper levels along transform systems. This suggests that fertile peridotites of the Coast Range ophiolite may be remnants that formed originally within a large- RFZ OFZ 1000 offset transform system. If this interpretation is correct, it has signifi cant

implications for subduction initiation. 1990) Kohler, and (Brey

1200

DISCUSSION T (°C)

Ophiolites and Subduction Initiation 1400 The progression from early MORB-like arc tholeiites (both as vol- 010203040 canic rocks and cumulate gabbros with forearc basalt parent magmas), 100 Cr/(Cr + Al) clinopyroxene

followed by boninitic lavas and related intrusions (wehrlite-pyroxenite CRO peridotites suite) represents the ﬁ rst two stages in the “life cycle” of suprasubduc- lherzolites ( ) tion-zone ophiolites recognized by Shervais (2001). This progression harzburgites ( ) Abyssal peridotites is characteristic of most suprasubduction-zone ophiolites and has long Small-offset or nontransform settings ( ) been correlated with forearc or intra-arc rifting (e.g., Pearce et al., 1981, Large-offset transform settings OFZ = Owen fracture zone (~300 km offset) 1984; Shervais, 1982; Shervais and Kimbrough, 1985; Juteau et al., VFZ = Vema fracture zone (~310 km offset) 1988; Crawford et al., 1989). The correlation of this progression with RFZ = Romanche fracture zone (~950 km offset) subduction initiation and nascent arc volcanism was recognized more Figure 8. Two-pyroxene equilibration temperatures (Brey and Köhler, 1990) recently (e.g., Shervais, 1990, 2001; Stern and Bloomer, 1992; Pearce, versus diopside Cr# for CRO (Coast Range ophiolite) peridotites compared 2003; Stern, 2004; Whattam and Stern, 2011). to abyssal peridotites associated with major fracture zones (Vema, Owen, The transition to calc-alkaline suite rocks (volcanic rocks and plu- Romanche) and normal, small offset transforms. Data sources: Hamlyn tonics of the TTD suite) in the Coast Range and other suprasubduction- and Bonatti (1980); Shibata and Thompson (1986); Dick (1989); Johnson et al. (1990); Bonatti et al. (1993); Edwards et al. (1996); Hellebrand et al. zone ophiolites is thought to represent the onset of stable subduction (2002); Brunelli et al. (2006). Mantle associated with large offset fracture (e.g., Shervais, 2001; Stern, 2004), but these rocks may represent in part zones is characterized by lower than ambient equilibration temperatures, evolved members of the preceding volcanic arc basalt-boninite suites most likely reﬂ ecting the hydrothermal cooling of the crust and mantle to (e.g., Shervais, 2008). This latter view is supported by the whole-rock deeper levels along transform systems.

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chemistry of the TTD suite—with relatively high MgO and Cr at high after subduction initiation, e.g., after subduction initiation along a frac- silica contents (Shervais, 2008)—and by their U-Pb zircon ages, which ture zone. Choi et al. (2008b) have proposed that this is the case, based are nearly coincident with subduction initiation at 172–166 Ma (Sher- on the low pyroxene equilibration temperatures that characterize these vais et al., 2005a). peridotites, and a comparison of these temperatures with fracture-zone The connection between ophiolites and subduction initiation has been and non-fracture-zone abyssal peridotites. These blocks must have been strengthened by the recent discovery of MORB-like forearc basalts as located suffi ciently far from the actual subduction interface to allow them a dominant volcanic assemblage in the Mariana forearc (Reagan et al., to persist and become mixed with the refractory peridotites, rather than be 2010). In the Marianas, this suite underlies the distinctive boninite suite removed by subduction erosion. and documents a period of nearly anhydrous decompression melting during the early stages of forearc spreading. The signifi cance of this discov- Early History of the Proto-Franciscan Convergent Margin ery is enormous, since it has commonly been assumed that the MORB- like early volcanic rocks of many ophiolites (and their cumulate gabbro Metasedimentary rocks of the Franciscan Eastern belt have Early to plutonic equivalents) represented true MORB formed at a mid-ocean- mid-Cretaceous depositional ages and mid- to Late Cretaceous meta- ridge spreading center, and this led to geodynamically suspect models of morphic ages (e.g., Blake et al., 1982; Wakabayashi and Dumitru, 2007; ridge-centered thrusting to initiate subduction (e.g., Oman; Boudier et al., Dumitru et al., 2010). These depositional and metamorphic ages are sig- 1988). These models were reinforced by the short elapsed time between nifi cantly younger than the overlying Coast Range ophiolite (Shervais et formation of the high-grade metamorphic soles and ophiolite formation al., 2005a; Hopson et al., 2008; Mattinson and Hopson, 2008). High-grade ages (e.g., Hacker, 1994). The recognition of slightly hydrous MORB- metamorphic blocks and rare coherent slabs with ages of 146–169 Ma like basalts that form in a forearc setting resolves these contradictions and (Lanphere et al., 1978; McDowell et al., 1984; Mattinson, 1986; Ross and fi nally permits a unifi ed model of subduction initiation that is consistent Sharp, 1988; Catlos and Sorensen, 2003; Anczkiewicz et al., 2004; Waka- both with ophiolite stratigraphy and with the volcanic evolution of in situ bayashi and Dumitru, 2007; Shervais et al., 2011) overlap with ages for forearc crust (Reagan et al., 2010). the Coast Range ophiolite and imply that the oldest blocks formed in the same nascent subduction zone as the ophiolite. Serpentinite Mélange and Subduction Initiation along a Transform Based on the oldest noninherited U-Pb zircon ages in the ophiolite (Kimbrough, in Shervais et al., 2005a; Hopson et al., 2008; Mattinson The connection between ophiolite formation and subduction initia- and Hopson, 2008), and the upper age range of high-grade metamorphic tion is now a robust and well-established paradigm. It is more diffi cult, blocks in the Franciscan assemblage (Ross and Sharp, 1988; Catlos and however, to document the exact nature of subduction initiation itself for Sorensen, 2003; Anczkiewicz, et al., 2004), we infer that proto-Fran- any specifi c example—Is it truly spontaneous in response to simple gravi- ciscan subduction initiated ca. 169–172 Ma (Aalenian-Bajocian). This tational forces, or induced by the compressional forces of plate dynam- age range shortly predates the oldest preserved radiolarian faunal assem- ics (e.g., Stern, 2004)? Does it nucleate on a fracture zone/transform, or blages in the ophiolite (Bajocian: Murchey, in Shervais et al., 2005a) and on some other zone of weakness? To answer these questions, we need to postdates compressive deformation in the Sierra Foothills by 5–20 m.y. examine the lithologic units that comprise the subduction assemblage and (see discussion in Shervais et al., 2005a). In contrast, Dumitru et al. the hanging wall of the subduction complex. (2010) documented deposition and metamorphism of the South Fork The former hanging wall of the proto-Franciscan subduction zone Mountain Schist at ca. 123 Ma using detrital zircon and white mica ages, appears to be preserved in the northern Coast Ranges as the Tehama- followed shortly by the Valentine Spring Formation and Yolla Bolly ter- Colusa mélange (Jayko and Blake, 1986; Hopson and Pessagno, 2005; rane. They concluded that this rapid onset of accretionary subduction Shervais et al., 2011). This serpentinite-matrix mélange contains elements followed a prolonged period of nonaccretionary subduction that lasted of the overlying forearc (suprasubduction) ophiolite complex, as well as some 40 m.y. (Dumitru et al., 2010). Many modern arcs are nonaccre- rocks derived from the underlying subducting oceanic plate (Shervais et tionary, and, in many cases, the forearc region has been subject to sub- al., 2011). Upper-plate (hanging-wall) rocks include volcanic and plutonic duction erosion (von Huene and Scholl, 1991; Scholl and von Huene, rocks derived from the ophiolite, sedimentary rocks from its overlying 2007). The signifi cance of this is that under these circumstances, frag- volcano-sedimentary cover sequence, and highly refractory peridotites ments of the original subduction interface may be preserved—albeit (harzburgites and dunites) that formed by hydrous melting in the man- highly deformed and mixed with younger rocks. tle wedge of the subduction zone (Choi et al., 2008a, 2008b; Jean et al., 2010). Lower-plate (footwall) rocks include the high-grade metamorphic THE MODEL blocks (amphibolite, garnet amphibolite, and high-grade blueschist with oceanic basalt affi nity) interpreted to have formed during subduction Spontaneous Subduction Initiation initiation, when the subducting oceanic crust was juxtaposed against hot hanging-wall peridotites (e.g., Wakabayashi, 1990). Fertile peridotites The most commonly applied model for spontaneous subduction initi- with compositions similar to abyssal peridotites are also found (Choi et ation involves sinking of older denser lithosphere along a transform-frac- al., 2008a, 2008b; Jean et al., 2010). While these have affi nities to the ture system (e.g., Karig, 1982; Casey and Dewey, 1984; Leitch, 1984; mantle lithosphere that underlies the footwall, their elevated fl uid-mobile Stern and Bloomer, 1992). Our model builds on these previous efforts element concentrations require prolonged residence within the upper and on later work (e.g., Bloomer et al., 1995; Shervais, 2001; Stern, plate, where hydrous fl ux from the subducting slab is rich in fl uid mobile 2004; Metcalf and Shervais, 2008). The initial condition is a large-offset elements (e.g., Jean et al., 2010). transform fault/fracture zone plate boundary with lithosphere of strongly Parkinson and Pearce (1998) have shown that fertile peridotites similar contrasting age on each side (Fig. 9A). As discussed by Stern (2004), to abyssal peridotites are commonly found within forearc regions, where spontaneous subduction initiation begins when the older lithosphere they coexist with highly refractory peridotites. The fertile peridotites may begins to sink gravitationally into the asthenosphere (Fig. 9B); this may represent former oceanic lithosphere that was trapped in the forearc region be induced in part by other factors affecting relative plate motions. As

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melts generated during this stage (“birth”) have MORB-like arc tholei- A oceanic crust ite chemistry, especially in the trace elements; silica extends to slightly higher concentrations in response to the limited fl uid infl ux (Fig. 9B). LM LM AM-2 In the lower crust, these MORB-like arc tholeiites form layered gabbros AM-1 Transform assemblage and gabbronorites. In the Coast Range ophiolite, these rocks include the early arc tholeiites and main layered gabbro series. Continued sinking of the dense lithospheric slab leads to cata- Transform SSZ forearc spreading B assemblage strophic dehydration of the low-temperature hydrous phases and fl ood- ing of the overlying mantle wedge with an aqueous fl ux that lowers the LM melting temperature of the already depleted (by stage-one decompres- Decompression sion melting) asthenosphere. The resulting melts are boninitic magmas AM-1 Melting: semidry with highly refractory major-element compositions and “U-shaped” or Stage 1: Forearc tholeiite AM-2 “spoon-shaped” trace-element concentrations (on mantle-normalized multi-element plots) that refl ect extreme melt depletion and fl uid-mobile C Transform SSZ forearc spreading element re-enrichment (e.g., Metcalf and Shervais, 2008; Whattam and assemblage Stern, 2011). Intrusion of the wehrlite-pyroxenite suite into the layered gabbro section refl ects this same hydrous melting event (Fig. 9C). Con- LM Slab dehydration AM-2 tinued extension of the nascent forearc crust during this stage leads to AM-1 Hydrous extensional plastic deformation of the layered gabbros, which enhances decompression intrusion of the less-deformed wehrlite-pyroxenite suite (e.g., Juteau et melting Stage 2: Boninite al., 1988; Hebert and Laurent, 1990; Shervais et al., 2004). The depleted harzburgites and dunites found in the mantle section of most ophiolites Transform refl ect this stage of melting. D assemblage Active island arc During this early extensional phase, the transform/fracture-zone LM Subduction assemblage on the upper plate forms the leading edge of the nascent erosion of forearc suprasubduction crust/lithosphere and remains adjacent to the plate AM-1 boundary as the lithosphere behind it extends over the sinking slab, and AM-2 Stage 3: Calc-alkaline the hinge line retreats away from the trench (Figs. 9B and 9C). This allows the transform/fracture-zone assemblage to remain relatively intact, even Figure 9. Schematic model for development of subduction initiation ophi- as the former oceanic crust behind it becomes subsumed by the newly olite, after Metcalf and Shervais (2008). (A) Juxtaposition of older, thicker formed forearc tholeiite/boninite crust (including their plutonic equiva- lithosphere on the west against younger, thinner lithosphere on the east; lents in the lower crust). The transform/fracture-zone assemblage is thus a transform assemblage of sheared lithospheric mantle and crust occu- preserved in a frontal position, trenchward of the forearc basin, where it pies the boundary between these zones. (B) Sinking of the older litho- can later interact with both upper-plate ophiolite assemblages and lower- sphere induces infl ow of asthenospheric mantle from below the younger plate ocean crust rocks. lithosphere; decompression melting of this lithosphere forms tholeiitic “forearc basalts” as new crust is created by the rapid extension above the Continued sinking of the slab eventually slows as a more stable sub- sinking lithosphere; rocks of the transform assemblage are carried west- duction geometry is achieved and hinge rollback slows (Fig. 9D). At this ward with the extending crust and mantle lithosphere. (C) As the older time, the plate boundary may become accretionary (adding new crust lithosphere continues to sink, dehydration reactions in the sinking slab through accretion of trench sediments) or nonaccretionary—where accre- release fl uids, which fl ux the overlying mantle wedge of asthenospheric tion is suppressed and erosion of the leading edge of the upper plate occurs mantle; the combination of continued decompression melting and water- (Scholl and von Huene, 2007). During this phase, the relict transform/ undersaturated melting of the asthenosphere (already partially depleted fracture-zone assemblage becomes broken up and mixed with upper-plate by prior decompression melting) results in boninitic magmas that erupt to form boninites sensu lato and intrude to form the wehrlite-pyroxenite lithologies (volcanics, plutonics, peridotites) and lower-plate lithologies series. Relict transform assemblage lithosphere continues to move at (mostly volcanics and sediments). The transform/fracture-zone assem- leading edge of extending upper plate. (D) The sinking slab stabilizes in blage also may be partially or wholly removed during this phase. position and confi guration, as balance between fl uid-fl ux melting and decompression melting obtains, leading to formation of a primitive island Subduction Erosion of the Forearc arc and eventually a transition to calc-alkaline composition lavas and intrusions. Subduction erosion of the leading edge of upper plate may be common during this phase. Black—oceanic crust; gray—lithospheric In most nascent arc systems, nonaccretionary or erosive subduction is mantle (LM); white—asthenospheric mantle (AM-1—beneath older plate; the rule, because the lack of an emergent arc limits the potential sediment AM-2—beneath younger plate). Stages refer to Shervais (2001). load available to the trench (Scholl and von Huene, 2007); this was apparently the situation during the early evolution of the proto-Franciscan subduction system (Dumitru et al., 2010). Over time, however, normal subduction leads to the formation of an emergent arc that can supply abundant shallow MORB-source asthenosphere fl ows upward to fi ll the void left sediment to the trench, which drives a transition from nonaccretionary by the sinking lithosphere, it undergoes decompression melting—anal- to accretionary subduction; this occurred in the proto-Franciscan subduc- ogous to the process at true mid-ocean ridges. These melts contain a tion system around 123 Ma (Dumitru et al., 2010). Continued evolution of minor amount of hydrous fl uids from the sinking slab but are otherwise the system could lead to arc rifting and the formation of a backarc basin, analogous to normal mid-ocean-ridge basalts (Stern and Bloomer, 1992; which would begin with suprasubduction-type magmas and evolve over Shervais, 2001; Metcalf and Shervais, 2008; Reagan et al., 2010). The time to a wider basin with MORB-like volcanics.

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Reconciliation of the Subduction Initiation Model with Evidence ACKNOWLEDGMENTS for Prior Continental Margin Subduction We wish to thank our colleagues for stimulating and insightful discus- One of the central conundrums of Cordilleran geology is the confl ict sions—including Robert Stern, Rodney Metcalf, Henry Dick, John between a subduction initiation model for the widely distributed Middle Wakabayashi, Sam Mukasa, and Marlon Jean—and cogent reviews by Jurassic ophiolites of California, Oregon, and Washington, and the occur- Jason Saleeby and Trevor Dumitru. This work was supported in part by rence of continental margin arc volcanism along a NW-SE–trending vol- National Science Foundation award EAR-44025 to Shervais. canic arc that began in the Early Triassic and continued into the Early Jurassic (e.g., Saleeby and Busby-Spera, 1992). Saleeby (2011) investi- REFERENCES CITED gated oceanic rocks of the Sierra Foothills belt and documented a long his- Alabaster, T., Pearce, J.A., and Malpas, J., 1982, The volcanic stratigraphy and petrogenesis of tory of ocean fl oor volcanism (Early Ordovician–Pennsylvanian) adjacent the Oman ophiolite complex: Contributions to Mineralogy and Petrology, v. 81, no. 3, to large-offset transforms that truncated the southwest margin of Laurentia p. 168–183, doi:10.1007/BF00371294. by sinistral shear. Subduction initiation occurred in the Late Permian along Anczkiewicz, R., Platt, J.P., Thirlwall, M.F., and Wakabayashi, J., 2004, Franciscan subduction off to a slow start: Evidence from high-precision Lu-Hf garnet ages on high grade- this truncated margin (Saleeby, 2011). 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Shervais, J.W., Choi, S.H., Sharp, W.D., Ross, J., Zoglman-Schuman, M., and Mukasa, S.B., 2011, Serpentinite matrix mélange: Implications of mixed provenance for mélange formation, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and MANUSCRIPT RECEIVED 10 MAY 2011 Societal Signiﬁ cance: Geological Society of America Special Paper 480, p. 1–30. REVISED MANUSCRIPT RECEIVED 25 AUGUST 2011 ° Shibata, T., and Thompson, G., 1986, Peridotites from the Mid-Atlantic Ridge at 43 N and their MANUSCRIPT ACCEPTED 12 OCTOBER 2011 petrogenetic relation to abyssal tholeiites: Contributions to Mineralogy and Petrology, v. 93, p. 144–159. Printed in the USA

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