Journal of Petrology Advance Access originally published online on June 3, 2005
Journal of Petrology 2005 46(10):2091-2128; doi:10.1093/petrology/egi048
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The Stonyford Volcanic Complex: a Forearc Seamount in the Northern California Coast Ranges
1 DEPARTMENT OF GEOLOGY, UTAH STATE UNIVERSITY, 4505 OLD MAIN HILL, LOGAN, UT 84322-4505, USA
2 DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF SOUTH CAROLINA, COLUMBIA, SC 29208, USA
3 DEPARTMENT OF GEOLOGICAL SCIENCES, SAN DIEGO STATE UNIVERSITY, SAN DIEGO, CA 92182-1020, USA
RECEIVED APRIL 15, 2004; ACCEPTED APRIL 6, 2005
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Jurassic age volcanic rocks of the Stonyford volcanic complex (SFVC) comprise three distinct petrological groups based on their whole-rock geochemistry: (1) oceanic tholeiites; (2) transitional alkali basalts and glasses; (3) high-Al, low-Ti tholeiites. Major and trace element, and Sr–Nd–Pb isotopic data indicate that the oceanic tholeiites formed as low-degree partial melts of normal mid-ocean ridge basalt (N-MORB)-source asthenosphere similar in isotope composition to the East Pacific Rise today; the alkalic lavas were derived from an enriched source similar to that of E-MORB. The high-Al, low-Ti lavas resemble second-stage melts of a depleted MORB-source asthenosphere that formed by melting spinel lherzolite at low pressures. Trace element systematics of the high-Al, low-Ti basalts show the influence of an enriched component, which overprints generally depleted trace element characteristics. Tectonic discrimination diagrams show that the oceanic tholeiite and alkali suites are similar to present-day basalts generated at mid-oceanic ridges. The high-Al, low-Ti suite resembles primitive arc basalts with an enriched, alkali basalt-like overprint. Isotopic data show the influence of recycled components in all three suites. The SFVC was constructed on a substrate of normal Coast Range ophiolite in an extensional forearc setting. The close juxtaposition of the MORB-like olivine tholeiites with alkali and high-Al, low-Ti basalts suggests derivation from a hybrid mantle source region that included MORB-source asthenosphere, enriched oceanic asthenosphere, and the depleted supra-subduction zone mantle wedge. We propose that the SFVC formed in response to collision of a mid-ocean ridge spreading center with the Coast Range ophiolite subduction zone. Formation of a slab window beneath the forearc during collision allowed the influx of ridge-derived magmas or the mantle source of these magmas. Continued melting of the previously depleted mantle wedge above the now defunct subduction zone produced strongly depleted high-Al, low-Ti basalts that were partially fertilized with enriched, alkali basalt-type melts and slab-derived fluids.
KEY WORDS: CRO; oceanic basalts; California
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Ophiolites have long been recognized as important elements of orogenic systems that indicate the formation and subsequent closure of oceanic basins (e.g. Moores, 1982
). They are especially important in regions such as the western Cordillera of North America, which comprises in part a tectonic collage of accreted terranes, as well as terranes formed by tectonic and magmatic processes in situ along the continental margin (Saleeby, 1992).
The Coast Range ophiolite (CRO) of California is one of the most extensive middle Jurassic ophiolite terranes in North America and has long been central to our understanding of Jurassic Cordilleran tectonics (e.g. Saleeby, 1992; Fig. 1). None the less, its origin is controversial and three primary hypotheses have been advanced: (1) formation at a mid-ocean ridge spreading center at low paleolatitudes, and its subsequent rapid drift northward to collide with North America (Hopson et al., 1981; Pessagno et al., 2000); (2) formation as a back-arc basin above an east-facing volcanic arc that collided with North America during the late Jurassic Nevadan orogeny (Dickinson et al., 1996; Godfrey & Klemperer, 1998; Ingersoll, 2000); (3) formation by forearc or intra-arc rifting along the western margin of North America, in response to nascent or renewed subduction of oceanic plates beneath North America (Shervais & Kimbrough, 1985a, 1985b; Stern & Bloomer, 1992; Shervais, 2001). It is clear that not all of these models can be correct, although there may be elements of truth in each. Most detailed studies of the CRO currently support the supra-subduction zone (SSZ) model (Shervais & Kimbrough, 1985a, 1985b; Stern & Bloomer, 1992; Shervais et al., 2004), but models based on back-arc basins and oceanic spreading centers still persist.
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The Stonyford volcanic complex (SFVC) is unique within the CRO (Fig. 2). This seamount complex consists almost entirely of volcanic flows (pillow lava, sheet flows) with subordinate diabase, hyaloclastite breccia, and minor sedimentary intercalations of chert and limestone (Shervais & Kimbrough, 1987
; Shervais & Hanan, 1989). Other normal components of the ‘classic’ ophiolite series (cumulate mafic and ultramafic rocks, isotropic gabbros and diorites, and sheeted dike complex) are found only as blocks in the serpentinite matrix mélange that underlies the volcanic complex. Most of the SFVC volcanic rocks show geochemical affinities to metavolcanic rocks of the Franciscan assemblage and they are distinctly different from the majority of CRO volcanics (Shervais & Kimbrough, 1987; Shervais & Hanan, 1989; Shervais, 1990; Shervais et al., 2004, 2005).
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In this paper we present field, petrological, geochemical, and isotopic data for rocks of the SFVC that establish its relationship to the CRO, and have implications for formation of the CRO and its subsequent tectonic evolution. These data also illuminate processes that may occur in active forearcs as a result of ridge–trench interactions.
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GEOLOGICAL RELATIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Overview
The Stonyford volcanic complex (SFVC) is located in the northern Coast Ranges of California approximately 120 miles north of San Francisco and 35 miles south of the Elder Creek ophiolite, near the western margin of the Sacramento Valley (Fig. 1). The SFVC crops out in portions of four USGS 7·5' quadrangles: Stonyford, St. Johns Mountain, Fout Springs, and Gilmore Peak. Mapping was conducted at a scale of 1:12 000 over two field seasons, covering
15 km2 (Fig. 2). The Stonyford area contains four main lithotectonic elements: (1) the SFVC; (2) low- to medium-grade schists of the Franciscan assemblage; (3) mudstone, greywacke, and conglomerate of the Great Valley Series; (4) a sheared serpentine-matrix mélange that separates metamorphic rocks of the Franciscan assemblage from unmetamorphosed sediments of the Great Valley Series, and encompasses the arcuate outcrop belt of the SFVC (Fig. 2).
Franciscan schists in this area include metagreywacke and argillite (quartz–albite–lawsonite schists), pale green metavolcanic rocks, and minor chert. These rocks, which correlate with the Yolla Bolly terrane of Blake et al. (1987), overlie chaotically sheared shale-matrix mélange of the Franciscan Central Belt to the north and SW. The contact between the Franciscan schists and the serpentinite-matrix mélange is generally a high-angle fault except along the northern border, where outcrop patterns imply a low-angle fault contact (Fig. 2).
The serpentine-matrix mélange may be a continuation of the Round Mountain mélange, which extends north to the Elder Creek remnant of Coast Range ophiolite (Jayko et al., 1987; Huot & Maury, 2002). The mélange matrix consists of strongly sheared and foliated serpentinite schist that crops out along ridge crests and in canyons (Shervais et al., 2005b). Tectonic knockers include massive harzburgite (up to 5 km x 2 km in size), unmetamorphosed sedimentary rocks (coarse lithic wackes, gritstones, and conglomerates—all rich in volcanic detritus), metasediments (foliated greywacke, argillite, gritstone), pale green metavolcanics, high-grade blocks of amphibolite and blueschist, fragments of the SFVC, and igneous rocks derived from the Coast Range ophiolite (wehrlite, clinopyroxenite, gabbro, diorite, quartz diorite, and keratophyre). These blocks of CRO plutonic and volcanic rocks are similar to those found near Elder Creek 60 km north of this area (Shervais, 2001; Fig. 1).
The Great Valley Series here is a homoclinal sequence of mudstone or argillite with minor intercalated greywacke and micritic limestone overlain by a coarse basaltic sandstone and the chert-rich Gravelly Ridge conglomerate (Brown, 1964). Buchia piochii of late Tithonian age are common throughout the section as high as the Gravelly Ridge conglomerate (Simpson-Seymore, 1999). The contact between the GVS and the serpentinite-matrix mélange is a high-angle reverse fault that dips steeply to the west and places serpentinite over sediments of the GVS. High-angle faults within the GVS that trend NW and west are probably tear-faults related to earlier west-vergent thrusting of the GVS (e.g. Wentworth et al., 1984; Glen, 1990).
Structural relationships
The SFVC is completely separated from metasediments of the Franciscan complex, and from mudstones and wackes of the Great Valley Series, by the serpentinite mélange that both underlies and overlies the massif. Nearly all contacts between the serpentinite mélange and the adjacent units are high-angle faults that reflect Neogene compression in the Coast Ranges.
The volcanic complex is divided into at least three large blocks by high-angle cross faults that trend east–west or SE–NW, and these blocks may be broken into smaller blocks along smaller faults (Fig. 2). One large, SE–NW-trending scissor fault has broken the largest block of volcanic complex just south of Auk Auk Ridge. Movement on this fault was normal, with the NW end of the upthrown block offset significantly more that the SE end, causing the upthrown block to tilt some 15–20° to the SE and exposing the low-angle contact between the serpentinite mélange and the SFVC on Auk Auk Ridge (Fig. 2). This fault and the relationships between the serpentinite mélange and SFVC are well exposed by deep erosion along the south side of Auk Auk Ridge, making this the best area in which to study contact relations between the volcanic complex and the mélange. S–C crenulation fabrics in the serpentinite matrix just west of Auk Auk Ridge and near Hyphus Creek to the south suggest normal shear sense on the low-angle contact with the volcanics (Dennis & Shervais, 1991).
The serpentinite-matrix mélange beneath Auk Auk Ridge contains a wide variety of ophiolitic rocks similar to those seen at Elder Creek (Shervais et al., 2004). These blocks include wehrlite, clinopyroxenite, gabbro, diorite, quartz diorite, and keratophyre pillow lava, all distinct from rocks exposed within the SFVC. We interpret these relations to suggest that the SFVC was built upon a substrate of ‘normal’ Coast Range ophiolite prior to (or perhaps during) its structural dismemberment and incorporation into the serpentinite mélange (see below).
Field relations within the SFVC
The SFVC consists largely of pillowed and massive lava flows with subordinate diabase, hyaloclastite breccia, and sedimentary intercalations (Shervais & Kimbrough, 1987; Shervais & Hanan, 1989; Zoglman, 1991). Lensoid intercalations of red radiolarian chert and pink siliceous mudstone up to 1 km long and 50 m thick occur throughout the section; limestone layers are smaller and less common. Although some folding is evident, most pillow lavas dip at moderate angles (40–60°) to the NE or east, with tops in the same direction. Field relations suggest that the volcanic complex represents a relatively intact submarine volcano that has been tilted to the NE and disrupted by high-angle faults.
Lavas
Lava flows in the SFVC include both pillow lava and sheet flows, but massive sheet flows seem to the dominant flow type. The base of the seamount sequence, which lies to the SW and west, is dominated by massive flows of oceanic tholeiites, with intercalations of pillow lava. Pillow lava becomes more common midway through the section and is commonly found intercalated with hyaloclastite breccia, which only occurs high in the section. The massive sheet flows range in thickness from about 1 m to over 4 m thick, and in places are seen to be laterally extensive. Pillows are typically 0·5–1 m across and form long tubes that branch and bud.
Hyaloclastite breccias
Hyaloclastite breccias form massive, unbedded layers which range from about 1 m thick to almost 100 m thick, although layers 10–50 m thick are most common. The hyaloclastite breccia layers are typically located near the top of the section, and presumably formed as the seamount shoaled. Three layers crop out in Dry Creek and four layers crop out between Dry Creek and Auk Auk Ridge (Glass, Fig. 2); other layers not shown in the map are less extensive and generally do not contain fresh glass.
All of the layers shown in the map contain fresh glass and were described by Shervais & Hanan (1989). These layers are intercalated with pillow lava, which allows us to determine tops. The breccias contain centimeter-scale rounded glass lapilli, along with 5–20% angular volcanic clasts up to 15 cm across in a matrix of glass shards and fragments. Pillow fragments, ‘finger pillow buds’, and small isolated (<
m) pillows are common in some breccias, but most are dominated by glass lapilli and shards. Small xenoliths of amphibolite are rare. These appear to represent fragments of the metamorphosed basement upon which the seamount was built. The hyaloclastite breccias were interpreted by Shervais & Hanan (1989) to represent submarine fire fountain deposits, based on the dominance of rounded glass lapilli and glass shards derived from round lapilli. Their occurrence near the top of the volcanic section implies shoaling of the volcano and a lowering seawater confining pressure on the magma.
Chert
Thick, extensive chert deposits occur throughout the SFVC; near the top of the section, pink siliceous mudstones are common. Limestone is scarce, forming small lenses <1 m thick. There are three thick chert lenses that are map-scale in size, which have been sampled extensively for biostratigraphic age determinations (Shervais et al., 2004). These are labeled Chert A, Chert B, and Chert C in Fig. 2. A fourth layer, Chert D, crops out at the famous ‘Diversion Dam’ locality (Fig. 2). This layer (and the intercalated pillow lava) has been extensively affected by hydrothermal fluids that jasperized the chert, forming massive ochre-colored jasper; radiolaria are recrystallized so that they can no longer be identified with confidence.
Chert A, which crops out just north of Stony Creek, is the thickest layer and also the most extensive, forming a mappable lens over 1 km long and up to 110 m thick, including basalt intercalations. The chert in this layer is manganiferous, and was exploited by a small Mn mining operation. Chert B crops out above Dry Creek, very close to the hyaloclastite breccias SFVG-1 to SFVG-4 (Glass 1–4, Fig. 2). This chert horizon is about 20 m thick and is well exposed along Black Diamond road. Chert C crops out along Black Diamond Creek, near the faulted margin of the complex, and in places it appears to be in fault contact with basalt underlying it. Chert C includes both ribbon chert and pink siliceous mudstone at least 4 m thick that contains rip-up clasts of the ribbon chert. Because of its faulted base a thickness could not be determined.
Age of the SFVC
Shervais et al. (2005a) presented age data on the CRO in northern California, including several dates for the SFVC and adjacent rocks. 40Ar–39Ar dates on volcanic glass from the hyaloclastite breccias range from 163 ± 0·7 to 164·7 ± 0·8 Ma, a narrow time interval corresponding to Bathonian on the time scale of Palfy et al. (2000). Cherts intercalated with the volcanic rocks contain radiolarian assemblages that imply a somewhat longer duration, ranging from Bajocian in the older cherts to Kimmeridgian or Tithonian in the youngest cherts. This corresponds to absolute ages ranging from c. 166–170 Ma near the base to c. 150–155 Ma near the top on the time scale of Palfy et al. (2000).
Quartz diorite mélange blocks that structurally underlie the SFVC yield U–Pb zircon concordia intercept ages of 163·5 ± 3·9 Ma and 164·8 ± 4·8 Ma (all uncertainties 2
). 207Pb/238U ages are similar to previously reported zircon ages for the CRO, ranging from 163 to 166 Ma (Shervais et al., 2005a). These ages are essentially the same as the volcanic glass and chert ages reported above, but may range to somewhat older ages. All of the age data reported for the SFVC by Shervais et al. (2005a) correspond to previously reported ages for the CRO, and cherts within the SFVC preserve the same fauna variations as observed in cherts of the CRO.
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PETROGRAPHY AND MINERALOGY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Volcanic rocks of the SFVC comprise three distinct petrological groups based on their petrographic characteristics and whole-rock geochemistry: (1) oceanic tholeiites with 1·9–3·3 wt % TiO2,
14 ppm Nb, Zr/Y <4, Ti/V = 27–32, and La/Smn = 0·67–0·9; (2) transitional alkali basalts and glasses with 1·7–2·9 wt % TiO2, 10–44 ppm Nb, Zr/Y = 2·4–7·8, Ti/V = 44–58, and La/Smn = 1·1–2·4; (3) high-Al, low-Ti tholeiites with <1·5 wt % TiO2, 16 ppm Nb (most <10 ppm Nb), Zr/Y = 0·8–7, Ti/V = 15–28, La/Smn = 0·34–1·8, and 15–24 wt % Al2O3. These distinctions are illustrated in Fig. 3, which compares their TiO2, and Al2O3, FeO*/MgO and Nb contents. Oceanic tholeiites form most of the complex and are found at all levels of exposure. Alkali basalts and basaltic glass are most common near the top of the seamount complex as interbedded pillow lava–hyaloclastite breccia sequences, but are also found beneath a thick chert lens low in the section, and commonly form pillows lower in the section. High-Al, low-Ti (HALT) lavas occur sporadically throughout the complex, and are also interbedded with hyaloclastites of alkali basalt. Some high-Al basalts are plagioclase megaphyric, but most are aphyric, so their high Al2O3 contents cannot be explained by simple plagioclase accumulation.
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Petrography
Oceanic tholeiites
Oceanic tholeiites are the most abundant of the three basalt suites. They crop out as black to reddish brown pillows and massive flows that form red to brick red soils. These basalts are predominantly aphyric or microphyric with equigranular to intergranular groundmass consisting of plagioclase, clinopyroxene, titanomagnetite, and ilmenite (Fig. 4a and b). Chemical modes calculated using the least-squares mixing program Genmix (Le Maitre, 1981
) show that average modal abundances of these rocks are 50% plagioclase, 38% clinopyroxene, 5·5% titanomagnetite, 3·6% olivine, and 2·3% ilmenite.
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Clinopyroxene forms colorless, non-pleochroic, granular to sub-ophitic grains of augite (Wo30–40 En36–53, mean Wo36 En46; mg-number 72) (Fig. 4). The pyroxenes are fresh and homogeneous, and range in size from 0·01 mm to 1·0 mm in diameter. Alteration of pyroxenes is to chlorite, magnetite, actinolite, and epidote. Plagioclase feldspar forms euhedral to subhedral laths, 0·05 mm to 1·25 mm long ranging in composition from An27 to An69; larger phenocrysts range from An61 to An91 (Fig. 6). Secondary mineral replacement may occur along the edges of the grains or along the twin planes within the crystal. Many samples contain euhedral feldspar with little or no alteration, whereas other samples show complete alteration and replacement of the feldspar by Ca-zeolites, albite, and prehnite. Pseudomorphs of olivine were identified in a few samples. They occur as microphenocrysts and represent <1% of the modal abundance of the sample. The grains are 0·05–0·3 mm in size, with complete replacement by chlorite and clay minerals. Titanomagnetite occurs as granular crystals ranging in size from 0·005 mm to 0·7 mm, found in the intergranular spaces between the plagioclase and pyroxene grains.
Alkali basalts
Alkali basalts are the second most common volcanic suite in the SFVC. They are most common towards the top of the volcanic sequence, but are found at all stratigraphic levels and underlie the lowermost chert lens in the seamount. The average chemical mode of the alkali basalts, calculated with Genmix, is 53% plagioclase, 29% clinopyroxene, 15% olivine, and 3% ilmenite. Other primary phases may include titanomagnetite and interstitial glass. Textures of the alkali basalts vary between sub-ophitic, porphyritic, and seriate, with up to 15% phenocrysts in the porphyritic samples (Fig. 4c and d).
Clinopyroxene in the alkali basalt series is typically titanian diopside (Wo31–46 En24–57, mean Wo43 En37; mg-number 66) with pale lavender color and faint pleochroism (Fig. 5). Granular pyroxenes range in size from 0·005 mm to 1·5 mm, sub-ophitic pyroxenes range in size from 0·005 mm to 2·5 mm. Plagioclase phenocrysts (An69–85) are subhedral to euhedral and range in size from 0·75 mm to 2 mm. The plagioclase phenocrysts may contain glass inclusions and are commonly altered, with albitic rims. The groundmass plagioclase (<0·75 mm long) is subhedral to anhedral, with compositions ranging from An24 to An70 where fresh, and <An20 where altered (Fig. 6). Alkali feldspar (Or20–93) is also found in groundmass of some samples. Titanomagnetite and ilmenite appear as minor primary phases in most samples. Secondary phases include chlorite, calcite, epidote, titanite, and clay minerals.
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Hyaloclastite breccias and volcanic glass
Alkali basalts erupted near the top of the volcanic edifice commonly form hyaloclastite eruption breccias that preserve unaltered volcanic glass (Shervais & Hanan, 1989
). Glass lapilli are contained in a matrix of smaller glass shards cemented by calcite, and altered in part to analcime and three compositional varieties of chlorite or interlayered smectite–chlorite (Shervais & Hanan, 1989). The glass lapilli consist of dark brown, transparent basaltic glass, with <2% modal microphenocrysts of olivine, plagioclase, and Cr-spinel; no clinopyroxene microphenocrysts were observed. All microphenocrysts are euhedral and show no indications of resorption or textural disequilibrium. Plagioclase (An72–80) and Cr-spinel remain relatively unaltered, but the olivine is typically replaced by smectites. Rare relict olivine is Fo85–87 in composition (Shervais & Hanan, 1989).
High-Al, low-Ti (HALT) basalts
These basalts include both plagioclase-megaphyric lavas that have accumulated feldspar and aphyric lavas with no feldspar accumulation. The Genmix-calculated chemical modes are 5–7% olivine, 60–77% plagioclase, 8–30% pyroxene, 5–7% titanomagnetite, and 0–1% ilmenite. Plagioclase phenocrysts compose anywhere from 2% to 30% of the mode in the porphyritic basalts. The groundmass of the plagioclase-phyric rocks is typically intergranular feldspar and pyroxene with one sample containing 10% devitrified glass. The aphyric samples exhibit textures ranging from equigranular to sub-ophitic and poikilitic.
Clinopyroxene forms subhedral granular to sub-ophitic grains of magnesian augite (Wo31–45 En38–65; mean Wo39 En50) with an average mg-number of 78 (Fig. 5), which range in size from 0·01 mm to 2 mm. Feldspar phenocrysts in the HALT suite range up to 5 mm across with compositions in the range An70–92·5 (Fig. 6). The less altered samples have euhedral plagioclase phenocrysts and lack zoning or albitic rims. More altered samples contain phenocrysts with albitic rims or anhedral grains completely altered to clay minerals. Some phenocrysts contain altered glass inclusions. The ragged edges of anhedral grains appear to have been resorbed as the remaining melt crystallized. Groundmass plagioclase (<0·5 mm long) comprises 50% of the groundmass in the porphyritic samples, with compositions that range from An31 to An79. In the aphyric basalts, plagioclase accounts for 65% of the mode with an average composition of An65. Plagioclase <An20 in composition replaces the rims of primary feldspar. Most samples contain either or both titanomagnetite and ilmenite, which form granular crystals <0·6 mm in size. Secondary phases include chlorite, albite, epidote, calcite, and clay minerals.
Mineral phase chemistry
Analytical methods
Mineral chemistry for major and minor elements was obtained with a Cameca SX50 four-wavelength-spectrometer automated electron microprobe at the University of South Carolina. Analyses were made at 15 kV accelerating voltage, 25 nA probe current, and counting times of 20–100 s, using natural and synthetic mineral standards. Analyses were corrected for instrumental drift, deadtime, and electron beam/matrix effects using the phi–rho–Z correction procedure provided with the Cameca microprobe automation system (Pouchou & Pichoir, 1991). Relative accuracy of the analyses, based upon comparison between measured and published compositions of the standards, is 1–2% for oxide concentrations >1 wt % and 10% for oxide concentrations <1 wt %. Mineral compositions are presented in Tables 1–3.
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Pyroxene
Clinopyroxene is ubiquitous in all volcanic rocks of the SFVC except for volcanic glass, where it is unknown (Table 1). Pyroxene is typically unaffected by low-temperature hydrothermal alteration and its compositions reflect primary magmatic compositions. Clinopyroxene varies from augite in the oceanic tholeiite series (Wo3040 En36–53) to titanian diopside or salite in the alkali basalt series (Wo31–46 En24–57); pyroxenes in the HALT series are magnesian augites (Wo31–45 En38–65). Pyroxene quadrilateral compositions are shown in Fig. 5.
Pyroxene minor element contents vary to some extent with both fractionation [mg-number = 100 x Mg/(Mg + Fe2+)] and magma series (Fig. 7). Pyroxenes in the HALT series are the most primitive, with mg-numbers ranging from 70 to 87 in most cases, and Cr2O3 contents as high as 1·5%. Pyroxenes of the oceanic tholeiite series and alkali basalt series have lower mg-numbers and generally lower Cr2O3, although some oceanic basalt Cpx has Cr2O3 as high as 0·8 wt % (Fig. 7). TiO2 and Na2O contents increase with decreasing mg-number in all series, and increase between series from HALT to oceanic tholeiite to alkali basalt. Alumina shows little correlation with either mg-number or magma series; some of this scatter may relate to cooling rate in groundmass pyroxene (e.g. Coish & Taylor, 1979).
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Crystal chemical substitutions among non-quadrilateral components are illustrated in Fig. 8. The high degree of correlation between Ti a.f.u. (atomic formula units, based on six oxygens) and tetrahedral AlIV implies that CaTi tschermakite (CaTiAlIV2O6) is the dominant coupled substitution (e.g. Schweitzer et al., 1979
). The offset of the data from the ideal ratio of Ti/AlIV = represents AlIV linked to other coupled substitutions (Fig. 8a). This offset largely disappears when CaCr tschermakite (CaCrAlIVSiO6) and CaFe3+ tschermakite (CaFe3+AlIVSiO6) are included (Fig. 8b); the remaining residual probably represents CaAl tschermakite substitutions (CaAlVIAlIVSiO6). Sodium shows no correlation with either Fe3+ or AlVI, except for a few alkali basalt pyroxenes that trend toward acmite component; only three analyses of alkali basalt pyroxenes show trends towards jadeite (Fig. 8c and d). Thus, acmite and jadeite are not important components in the oceanic tholeiite or HALT magma series, and form minor components in the alkali basalt series. This suggests that sodium must be charge balanced by other coupled subsititutions, such as NaTiAlIVSiO6, as proposed by Schweitzer et al. (1979). It is also consistent with the observed absence of blueschist-facies mineral phases in any of the SFVC rocks studied to date.
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Feldspars
Feldspar is the most abundant of the major phases, with modes ranging from 50% to 75%. Primary feldspar compositions are commonly preserved, but feldspars are more likely than pyroxene to be altered to lower-temperature assemblages. Plagioclase phenocrysts in all series range from
An60 to An92 whereas groundmass plagioclase is typically more sodic, ranging from An24 to An70 (An79 in the HALT series; Fig. 6; Table 2). Groundmass plagioclase and the rims of phenocrysts are commonly ‘albitized’, with clear plagioclase <An20 mantling cloudy cores. Alkali feldspar (Or20–93) is rare, being found only in the groundmass of some alkali basalts or as intergrowths within plagioclase.
Plagioclase phenocrysts with compositions more calcic than An80 are found in all magma series (Fig. 6). These phenocrysts appear to be out of equilibrium with their host magmas, and may represent relict phenocrysts preserved from a more primitive magma that was mixed with a more evolved magma prior to eruption. Such mixing events are common in oceanic magma systems (e.g. Stakes et al., 1984).
Olivine
Olivine is no longer present in the basalts, but pseudomorphs of chlorite and clays after olivine are found. Relict olivine of Fo85–87 composition is found within some glass lapilli (Shervais & Hanan, 1989), but it is rare. These olivines form euhedral prisms that are in equilibrium with the host glass and are clearly cognate phenocrysts (Table 1).
Oxides
The predominant accessory oxide phase is titanomagnetite with TiO2 ranging from 3 to 29 wt %. MnO ranges from 1 to 7 wt % and correlates positively with TiO2 and inversely with FeO. Titanomagnetite is most abundant within the oceanic tholeiite suite and occurs throughout the crystallization sequence. Titanomagnetite appears at the very end of crystallization in the high-Al, low-Ti basalts. Ilmenite within the alkali basalts occurs as an early crystallizing phase, whereas titanomagnetite appears in the later stages of crystallization. Ilmenite contains 2·8–4·7% MnO.
Cr-spinel is rare in the volcanic rocks studied here, but relatively common in the volcanic glasses (Table 3). These spinels are all titaniferous magnesiochromites, which indicates alkali magmatic affinities (e.g. Sigurdsson & Schilling, 1976). One spinel analyzed from an alkali basalt is also a titaniferous magnesiochromite. One spinel was analyzed from a high-Al, low-Ti basalt and found to be a chromian spinel. Spinel was not found in oceanic tholeiite series.
Secondary phases
Secondary phases formed in response to low-temperature metamorphism include chlorite, albite, titanite, quartz, epidote, prehnite, pumpellyite, zeolites, smectites, and calcite; actinolite is rare. Chlorites are aluminous, similar to chlorites in Franciscan metabasalts, and distinct from quartz–chlorite breccias and metavolcanics of the Mid-Atlantic Ridge. Olivines are commonly altered, but pyroxene and plagioclase are typically preserved. As a result, whole-rock geochemistry is only marginally affected by low-temperature metamorphism.
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WHOLE-ROCK CHEMISTRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Analytical methods
Whole-rock major element analyses were determined by a fused glass bead technique for 10 major elements [SiO2, TiO2, Al2O3, FeO* (total iron as FeO), MnO, MgO, CaO, Na2O, K2O, and P2O5] using the Cameca SX-50 electron microprobe at the University of South Carolina and natural glass standards (Shervais et al., 1990
). Whole-rock powder was fused in an electric induction furnace with a nitrogen atmosphere using molybdenum foil boats; the resulting glass chips were mounted on glass microprobe mounts and polished prior to analysis. The fused glass beads were analyzed using the same conditions as for mineral analysis and standards selected for glass analysis. The natural Juan de Fuca glass VG2 was used for most elements, supplemented by mineral standards for minor elements. Na was analyzed first to minimize effects of diffusion. Repeated analyses of USGS whole-rock standards show that analytical uncertainty is <2% relative for major elements and <5–10% relative for minor elements (i.e. elements with concentrations of <1 wt %; Shervais et al., 1990).
Selected trace elements (Nb, Zr, Y, Sr, Rb, Zn, Cu, Ni, Cr, V, Sc, Ba) were analyzed by X-ray fluorescence (XRF) spectrometry using a Phillips PW 1400 XRF system equipped with a Rh tube. Powdered sample (6–8 g) was combined with 4–6 drops of a 2% polyvinyl alcohol solution, pressed into pellets, and dried overnight at 110°C. Calibration was based on 12 well-characterized USGS rock standards. Additional trace elements [Th, Hf, Ta, Co, and the rare earth elements (REE)] were determined by instrumental neutron activation analysis (INAA) for 40 samples at the Oregon State University Radiation Center. Samples were irradiated in the OSU TRIGA reactor and counted following the procedures outlined by Laul (1986). Analytical uncertainties are typically 2–5% relative for most elements, except Nd (10%). Major and trace element data are presented in Table 4.
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Major elements
Oceanic tholeiites of the SFVC are characterized by high FeO* (>12 wt %) and TiO2 concentrations (>2 wt %) that increase with decreasing MgO, and by low CaO (
8–12 wt %) and Al2O3 (12–16 wt %) concentrations that decrease with decreasing MgO (Fig. 9). The alkali basalts and glasses have TiO2 concentrations similar to the oceanic tholeiites, but they are lower in FeO* and higher in Al2O3, all of which increase or remain constant with decreasing MgO (Fig. 9; all data recalculated anhydrous).
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The HALT basalts are characterized by high Al2O3 (
15–24 wt %) and CaO concentrations (10–18 wt %), and by low TiO2 (<1·5 wt %) and FeO* (10 wt %) concentrations that decrease with decreasing MgO (Fig. 9). CaO concentrations increase, and alumina concentrations are relatively constant with fractionation. Because they have low FeO* at relatively high MgO, FeO*/MgO ratios in the HALT suite are low for a given MgO concentration, in contrast to the oceanic tholeiite suite, which extend to higher FeO*/MgO ratios (Fig. 3). This implies formation of the HALT suite by partial melting of a source region that was more refractory than that sampled by the tholeiitic suite. The extremely high concentrations of Al2O3 (>19 wt % in some samples) may be attributed to plagioclase accumulation, but the overall high Al2O3 concentrations in the aphyric and sparsely phyric samples must represent the actual liquid composition.
The alkali elements, Na2O + K2O, range from 2 wt % to as high as 7 wt %, with most of the higher concentrations being attributed to low-temperature hydrothermal alteration of the primary basalts. However, the volcanic glasses show no such effect, indicating that their concentrations of 3–4 wt % alkalis are primary and probably represent the highest values expected in the unaltered protoliths. Two alkali basalts with the highest total alkali contents (5·4–7·0 wt %) also have extremely low CaO (<6·2 wt %) and Cr (100 ppm), and the highest observed MgO (8·6 wt %) concentrations (Fig. 9); one sample also has extremely high Ba (6400 ppm). This combination of contrasting refractory and evolved characteristics suggests that these two samples have been seriously compromised by low-temperature hydrothermal alteration, including the exchange of Ca in the protolith for Mg and alkalis in seawater. Most samples do not show these effects and may be used to reconstruct primary magma evolution.
Trace elements
The oceanic tholeiite and alkali suites are generally higher in high field strength elements (HFSE; e.g. Zr, Nb) and lower in the plagiophile elements (Sr, Rb) and refractory elements (Cr, Ni). In contrast, the HALT suite is characterized by higher concentrations of refractory and plagiophile elements, and lower HFSE (Fig. 3). In detail, the alkali glasses have higher Cr and Ni (comparable with the HALT basalts) than the alkali basalts, consistent with their high MgO concentrations, and all of the volcanic suites have variable Rb concentrations, consistent with their variable alkali concentrations and modest low-temperature mobilization of these elements. Nb concentrations are generally much higher in the alkali basalts and glasses (10–44 ppm) than in comparable oceanic tholeiites (14 ppm) and HALT basalts (10 ppm).
Sc concentrations are highest in the oceanic tholeiite suite (34–55 ppm, with most >40 ppm), suggesting derivation from a source rich in calcic pyroxene. Sc concentrations are lower in the alkali basalts (Sc 17–35 ppm) and glasses (Sc 28–31 ppm), suggesting that they were formed by larger degrees of partial melting that eliminated calcic pyroxene in their source residues. The HALT basalts also have relatively low Sc concentrations (25–38 ppm) that imply either higher degrees of partial melting or derivation from a source in which a previous melting event removed the low melting components such as calcic pyroxene and the trace elements they contained. This latter hypothesis is consistent with the high refractory element contents (Mg, Cr, Ni) of the HALT series.
The REE concentrations are shown in Fig. 10 normalized to the average C1 chondrite of Sun & McDonough (1989). Oceanic tholeiites of the SFVC are characterized by light REE (LREE)-depleted patterns (La/Smn = 0·67–0·9) that parallel normal MORB (N-MORB; Sun & McDonough, 1989) but with concentrations two to three times higher than N-MORB (La 16–24 times chondrite). Patterns for the middle REE (MREE) and heavy REE (HREE) are relatively flat, with a slight negative Eu anomaly (Fig. 10a). Alkali basalts and glasses are enriched in the LREE (La/Smn = 1·1–2·4) and have patterns that slope continuously from La to Lu (Fig. 10b). They are more enriched than the average E-MORB (enriched MORB) of Sun & McDonough (1989) but less enriched than the average ocean island basalt (OIB), with La concentrations 5·5–9·8 times N-MORB. The HALT basalts have a wide range of REE patterns, from LREE-depleted to LREE-enriched (La/Smn = 0·34–1·8), and La concentrations that range from 0·3 to 3·3 times N-MORB (Fig. 10c). The REE patterns of these samples are distinct from those of both the more common oceanic tholeiite series and the alkali basalt–glass series, and cannot be related to either of these series by any simple magmatic processes.
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Multi-element variation diagrams of incompatible and compatible elements normalized to the average N-MORB of Sun & McDonough (1989)
are presented in Fig. 11. Elements are arranged in this diagram from those most incompatible in refractory mantle assemblages on the left to those most compatible with refractory mantle assemblages on the right (Thompson et al., 1983; Sun & McDonough, 1989). Oceanic tholeiites are characterized by abundance patterns that parallel N-MORB for the less incompatible elements (La to Na), are relatively depleted in the more compatible elements (Fe to Ni), and highly enriched in the most incompatible elements (Rb to K; Fig. 11a). Phosphorus is depleted in a few samples, all of which were affected by strong saprolitic alteration. Abundance patterns of the SFVC ocean island tholeiite suite suggest derivation from a mantle source dominated by MORB-source asthenosphere, somewhat enriched by a low-melt component (high Nb, Ta) and possibly a subduction component (high Rb, Ba), although this may reflect later hydrothermal alteration.
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Alkali basalts and glasses are characterized by abundance patterns that are enriched relative to N-MORB for all elements more incompatible than Y; they are depleted in the more compatible elements relative to N-MORB (Fig. 11b). A few samples show positive spikes in Rb, Ba, K, and Sr that may be attributed to low-temperature alteration, but most display smooth patterns that decrease systematically from Rb to Sc, consistent with derivation from an enriched source similar to that proposed for E-MORB and OIB. This source is presumably a mixture of N-MORB source asthenosphere with plume source mantle, similar to that observed along the Reykjanes Ridge south of Iceland (Sun et al., 1975
).
The HALT basalts are characterized by abundance patterns for the less incompatible (Zr to Na) and compatible (Zn to Ni) elements that are parallel to or depleted relative to N-MORB, with the more incompatible elements ranging from depleted to strongly enriched relative to N-MORB (Fig. 11c). These patterns are similar to those observed in forearc ophiolite lavas, with strong enrichment in the lithophile low field strength elements superimposed on less incompatible REE and high field strength elements that are more depleted than N-MORB (Metcalf et al., 2000). In contrast to this simple scenario, however, the HALT basalts are also enriched in the more incompatible ‘high field strength’ elements Nb and Ta as well, suggesting source enrichment similar to that observed in the alkali suite.
The nature of the enrichments seen in the MORB-normalized trace element diagrams can be examined using incompatible element ratio plots in which each ratio has the same denominator, so that mixing of sources results in a linear array, and fractional crystallization has no effect. A plot of Zr/Nb vs Y/Nb (Fig. 12) shows mixing between an enriched OIB or hotspot-source mantle (low Zr/Nb and Y/Nb) and depleted MORB-source mantle (high Zr/Nb and Y/Nb). The SFVC alkali basalt series plots near the enriched source composition, whereas the SFVC oceanic tholeiite series forms an array that ranges from somewhat higher than plume-like source ratios to MORB-like ratios (Fig. 12). In contrast, the HALT series forms an array that ranges from low ratios, similar to the alkali basalts, to ratios that are much higher than MORB. These super-MORB ratios must reflect melting of a source that was depleted in Nb relative to Zr and Y, prior to the addition of an enriched, plume-like component.
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ISOTOPE GEOCHEMISTRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Analytical methods
Thirteen whole-rock and glass samples were selected for Pb, Sr and Nd isotope analysis and for determination of the Pb, U and Th elemental abundances following methods similar to those described by Hanan & Schilling (1989)
and Graham et al. (1998). The Pb, Sr and Nd isotopes and the Pb, U and Th elemental abundances are presented in Table 5. The measured Sr, Nd, and Pb isotope compositions have been adjusted for radiogenic decay to their initial isotope ratios at 164 Ma (Shervais et al., 2005a) using the XRF and INAA data (Table 4) for Rb, Sr, Nd and Sm and isotope dilution data for U, Th and Pb elemental abundances (Table 5). Co-variation diagrams for 143Nd/144Nd164 Ma–87Sr/86Sr164 Ma, 143Nd/144Nd164 Ma–206Pb/204Pb164 Ma, and 87Sr/86Sr164 Ma–206Pb/204Pb164 Ma are shown in Fig. 13; co-variation diagrams for the Pb isotopes are shown in Fig. 14.
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The volcanic glasses are free from visible alteration. All of the basalts have been affected to some extent by low-temperature alteration at sub-greenschist-facies conditions, but the extent of this alteration is generally minor. Olivine is typically replaced by chlorite–smectite and plagioclase may have albitic rims, but in general plagioclase, pyroxene, and oxides all retain their primary compositions. All samples were carefully handpicked, to avoid alteration, secondary vein minerals, and vesicle fillings, and then rinsed in dilute acids to remove pollution that may have taken place during preparation.
Results
Basalts of the oceanic tholeiite suite have slightly more depleted Nd, Sr, and Pb isotope signatures relative to the alkali basalts and HALT lavas. The 143Nd/144Nd164 Ma (0·512899–0·512906) and 87Sr/86Sr164 Ma (0·70232–0·70296) initial ratios of the oceanic tholeiite samples show little variation, and fall within the range of Nd and Sr initial isotope ratios reported for Mesozoic (130–151 Ma) MORB from Deep Sea Drilling Project (DSDP) sites in the western Pacific (Janney & Castillo, 1997). The alkali basalts have 143Nd/144Nd164 Ma (0·512268) and 87Sr/86Sr164 Ma (0·70337–0·70394) initial ratios that are enriched relative to the western Pacific Jurassic MORB and to the modern, zero-age East Pacific Ridge (EPR) MORB field (Hanan & Schilling, 1989; Janney & Castillo, 1997; Fig. 13a). The HALT sample SFV 37-1 has an unusually high Rb/Sr ratio of 0·23 and very low 87Sr/86Sr164 Ma of 0·70294.
The Pb isotope initial ratios are more radiogenic than the 130–151 Ma Pacific MORB reported by Janney & Castillo (1997) except for two 132 Ma DSDP leg 17, site 166 basalts that have similar 206Pb/204Pb164 Ma initial ratios (Fig. 14). Whereas the SFV lavas have 206Pb/204Pb164 Ma ratios within the range of the zero-age EPR and Easter microplate region MORB, the 207Pb/204Pb164 Ma and 208Pb/204Pb164 Ma ratios are relatively higher for the 206Pb/204Pb164 Ma, and trend toward and overlap with the field for Pacific ocean-floor sediment (Fig. 14). The trend toward oceanic sediment is also apparent in the 206Pb/204Pb164 Ma vs 87Sr/86Sr164 Ma and 143Nd/144Nd164 Ma isotope variation diagrams (Fig. 13b and c). The lack of significant alteration and the tendency for the alkali basalts to fall out of the range for Pacific MORB, toward the field of oceanic sediment, in Nd–Sr–Pb isotope multi-isotope diagrams are consistent with a recycled sediment component in the mantle source.
The measured 238U/204Pb ratios are 21–35 in OIB, 12–35 in alkalic basalt whole rocks, and 112 in the HALT basalt. These ranges in 238U/204Pb are elevated relative to the 12·7 ± 5·78 range reported for unaltered Pacific MORB (White, 1993). The measured 232Th/238U ratios in the whole rocks are 1·3–2·4 in OIB, 2·9–4·2 in alkalic basalt whole rocks, and 0·17–2·8 in the HALT basalt, compared with 2·7± 0·345 in fresh Pacific MORB (White, 1993). The lack of significant alteration in the basalts suggests that the high 238U/204Pb, relative to MORB, is a source feature and not due to sea-floor alteration-related U uptake. However, the very low 232Th/238U and high 238U/204Pb of the HALT basalt SFV-37-1 is suspect. When this basalt is age corrected its Pb and Sr isotope compositions appear to be overcorrected for U and Rb decay in the 206Pb/204Pb and 87Sr/86Sr vs 206Pb/204Pb diagrams (Fig. 13). The limited alteration of the HALT basalt argues against seawater alteration as an explanation, and suggests that assimilation of a sea-floor sediment (± seawater) component may have been a factor.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
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Petrogenesis
Major element fractionation modeling
Each of the volcanic suites studied here displays systematic linear trends in elemental variation diagrams that may result from fractional crystallization of more primitive parental magmas. Least-squares mixing models were used to calculate the extent of crystal fractionation within each volcanic suite, using the program Genmix (Le Maitre, 1981). Fractionation was determined by mixing average phenocrysts compositions with an evolved magma composition and determining the proportions of each required to create the most primitive sample composition.
The oceanic tholeiite suite ranges in MgO from 8 wt % to 3·2 wt %, with FeO*/MgO ratios of 1·3–3·3 (Fig. 9). Using sample SFV-91-1 (7·98 wt % MgO) as the parent magma, 71% fractional crystallization is needed to create the evolved sample SFV-114-1 (5·0 wt % MgO). The fractionating phase assemblage is plagioclase (52%), pyroxene (37%), olivine (10%), and ilmenite (1·4%). The sum of the square of the residuals is 0·15, suggesting a robust fit to the data. The fractionating assemblage is close to the average chemical mode for this suite (plagioclase 50%, pyroxene 38%, olivine 4%, and oxides 8%), but with more fractionating olivine and less fractionating oxides, which supports the conclusion that this calculation provides a good approximation of the actual fractionating assemblage.
The alkali basalts and glasses range in MgO from 7·8 to 2·2 wt %, with FeO*/MgO ratios of 1·0–3·2, although most samples have MgO >4·5 wt % (Fig. 9). Using glass SFVG-5 (7·28 wt % MgO) as the parent magma, 40% fractional crystallization is needed to create the evolved sample SFVP-1 (4·9 wt % MgO). The fractionating phase assemblage is plagioclase (52%), pyroxene (31%), olivine (16%), and ilmenite (1%). The sum of the square of the residuals is 0·73, suggesting a reasonable fit to the data. The fractionating assemblage is close to the average chemical mode for this suite (plagioclase 53%, pyroxene 29%, olivine 15%, and oxides 3%). The high proportion of olivine in the fractionating assemblage, compared with the ocean island tholeiite suite and the HALT suite, explains why the alkali basalts require almost half the fractional crystallization of the other suites to reduce MgO concentration by the same amount as seen in the other suites.
The high-Al, low-Ti basalts range in MgO from 7·6 to 4·2 wt %, with FeO*/MgO ratios of 1·0–1·9 (Fig. 9). Using sample SFV-123-1 (7·66 wt % MgO) as the parent magma, 83% fractional crystallization is needed to create the evolved sample SFV-Z47-1 (4·5 wt % MgO). The fractionating phase assemblage is plagioclase (65%), pyroxene (20%), olivine (8%), titanomagnetite (6%), and ilmenite (1%). The sum of the square of the residuals is 1·35, suggesting a poor fit to the data. The fractionating assemblage is close to the average chemical mode for this suite (plagioclase 68%, pyroxene 21%, olivine 6%, and oxides 7%). The proportion of plagioclase in the fractionation assemblage of the HALT suite is consistent with its high Al2O3 and CaO concentrations, and low concentrations of mafic components (FeO*, TiO2). The poor fits obtained for this suite imply that these samples are not related by simple fractional crystallization alone.
In summary, each of the magmatic suites studied here displays major element chemical variations that are consistent with fractional crystallization of distinct phenocryst assemblages from parent magmas that are also distinct in their chemical compositions. There is no indication that any of these suites can be related to one another by any simple magmatic processes, a conclusion that is supported by their trace element systematics as well. We conclude that although these magma suites may share a common or related mantle source region, they have remained largely independent of one another since separation from their mantle source regions. This requires that each magma series exploited its own plumbing system to reach the surface, and, because these suites are intercalated at all levels of the volcano, that these plumbing systems may have been active concurrently. Similar conclusions have been reached for the plumbing systems of large Hawaiian volcanoes, which also erupt distinct magmas that show no signs of chemical commingling (e.g. Wright, 1971; Tilling et al., 1987).
Trace element modeling
Incompatible trace elements in the SFVC volcanic suites exhibit limited enrichment in response to fractionation (decreasing MgO), with maximum fractionation factors of 40–60% fractional crystallization. This compares well with fractionation factors based on major element modeling of 40–80% fractional crystallization (see above). In detail, the enrichment in trace elements between potential parent–daughter pairs commonly implies different amounts of fractionation than the major element models, with the major elements indicating less fractionation than the trace elements. This suggests that much of the variation in trace element concentrations results from differences in the parent magmas within each suite, and is not controlled exclusively by fractional crystallization.
Differences within each suite in incompatible trace element concentrations must be due in part to different degrees of partial melting of similar source regions. In contrast, differences between the three suites defined here (oceanic tholeiite, alkali basalt, high-Al/low-Ti) probably represent partial melting of distinct source regions that differ significantly in their trace element characteristics. To model these variations, we have constructed a series of melting models featuring spinel or garnet lherzolite modes, and chemical compositions that reflect the source regions of N-MORB, E-MORB, and a depleted N-MORB source. Partition coefficients were taken from Arth (1976) and McKenzie & O'Nions (1991, 1995). The N-MORB source composition was taken from McKenzie & O'Nions (1995), the E-MORB source composition was taken from Mertz et al. (2001), calculated as the N-MORB source + 8% metasomatic melt (0·3% fractional melt of MORB source; Mertz et al., 2001). The depleted N-MORB source was calculated by subtracting a 1% melt fraction from the N-MORB source composition of McKenzie & O'Nions (1995). Melting calculations were carried out using the non-modal batch melting equation; results are presented in Fig. 15 normalized to the primitive mantle composition of McKenzie & O'Nions (1995).
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The oceanic tholeiite suite provides good fits at 3–5% partial melting of a spinel lherzolite, N-MORB source (Fig. 15a). The LREE are slightly depleted relative to the HREE, and the HFSE Ti, Zr, Hf, Nb, and Ta are enriched relative to the HREE. Other models (not shown) show that the oceanic tholeiite suite cannot be formed by melting an E-MORB type source, regardless of the mantle mode used. In addition, the alkali basalt–glass suite cannot be derived from melting an N-MORB composition source: the LREE/HREE ratios of the model melts reflect those of the source (that is, LREE-depleted) except for melt fractions <1%, but even here the LREE/HREE ratio of the model melt is too low to match the alkali basalt–glass suite; other elements show equally poor fits (Fig. 15b). The alkali basalt–glass suite is well matched by models that melt an E-MORB composition source at relatively shallow depths (spinel lherzolite modes). Good fits are obtained for all elements at 10–15% melting; a more enriched source would result in smaller melt fractions (Fig. 15c). Models that use the same source composition but with a garnet lherzolite mode show poor fits with data: LREE/HREE ratios are too high in the model melts, resulting in REE patterns that cross the alkali basalt–glass REE patterns, and HFSE contents that are too high (Fig. 15d).
The HALT suite shows a wide range of incompatible trace element concentrations that cannot be fitted by any simple models. The most depleted HALT samples cannot be modeled as melts of spinel lherzolite N-MORB source at <20% partial melting (Fig. 15e). More enriched samples match the model REE patterns approximately at 3–5% melting, but results for K, Sr, and the HFSE are poor. Melting of a depleted N-MORB source works somewhat better: REE patterns match approximately at 7–15% melting, but results for K, Sr, and the HFSE are still poor (Fig. 15f). The highly depleted nature of the HALT source is supported by the Nb/Zr/Y systematics (Fig. 12) and by their high mg-number and Cr content; these data also suggest that the depleted melts may have mixed with an enriched component similar to the alkali basalt–glass suite to achieve the observed range in incompatible element concentrations. If we use the depleted HALT samples as potential ‘primary melts’, we can reproduce the trace element concentrations of the more enriched HALT samples by mixing the ‘primary melts’ with up to 50% alkali basalt–glass melts. This mixing model provides the best fit to the enriched HALT melts, and supports inferences drawn from their major and trace element systematics for enrichment of a previously depleted source region by an enriched OIB melt component. The radiogenic isotope signature of the alkali basalts requires a recycled sediment component in the mantle source.
In summary, the oceanic tholeiite suite represents low-degree partial melts (3–5%) of an N-MORB asthenosphere source region at relatively shallow levels (spinel lherzolite facies). The alkali basalt–glass suite represents moderate degrees of melting (10–15%) of an E-MORB–OIB-like source that was physically distinct from the oceanic tholeiite source. Finally, the HALT suite appears to represent second-stage melts from a depleted MORB source that have been variably enriched by melts similar to the alkali basalt suite.
Metamorphism in the SFVC
Replacement of minerals in basalt
Metamorphic assemblages found in the SFVC are typical of low-temperature ocean-floor metamorphism. Metamorphic phases include chlorite, albite, epidote, prehnite, calcite, titanite, Ca-zeolites, clay minerals, and quartz; actinolite also occurs in some samples. Plagioclase laths may be albitized, especially along their margins, and cores may be replaced by epidote, prehnite, calcite, Ca-zeolites, or clay minerals. Albitization is commonly associated with the alteration of primary magnetite or ilmenite to magnetite–titanite intergrowths. Olivine phenocrysts are rare and when found are totally replaced by green chlorite and smectite, except in the volcanic glass. The pyroxenes remain relatively fresh and show alteration only along grain edges and fractures. The pyroxenes alter to chlorite, actinolite, epidote, and titanite.
Despite the widespread occurrence of secondary minerals in volcanic rocks of the SFVC, preservation of primary minerals is extensive and shows that circulation of hydrothermal fluids was not pervasive and that water/rock ratios were low. The most extensive alteration of SFVC volcanic rocks is found along Stony Creek in the western part of the complex. Here, large tracts of pillow lava have been extensively replaced by clay minerals, calcite, and chlorite (‘brownstone’). Chert horizons stratigraphically above this area of extensive clay-alteration are jasperized to massive, red and ochre-colored chalcedony. We interpret these relations to document the location of an extensive submarine hydrothermal vent system that was active prior to disruption and emplacement of the complex.
Hydrothermal alteration of the hyaloclastite breccias
Secondary mineral assemblages in the hyaloclastite breccias include calcite, analcime, chlorite, and chlorite–smectite intergrowths (Shervais & Hanan, 1989). Initial alteration began with replacement of some glass lapilli by smectite–chlorite intergrowths. Calcite cementation possibly deposited by cold, circulating seawater sealed off the stratigraphically lower horizons. The calcite cement virtually blocked off penetration of hydrothermal fluids coming from depth or from seawater circulating down from the surface. After filling of remaining voids by chlorite, calcite, and analcime, the last stage of metamorphism involved further replacement of glass lapilli by Ca-zeolites. Textures of the secondary minerals show that deposition of these phases occurred early, prior to extensive alteration of the volcanic glass and preserving the glass from extensive replacement (Shervais & Hanan, 1989).
Vein minerals and vesicle fillings
Vein assemblages and vesicle fillings are common in the SFVC. Calcite, prehnite, and Ca-zeolites (thomsonite, laumontite) constitute the most common vein fillings; quartz, analcime, and epidote are less common. Larger veins are commonly lined by Ca-zeolites along their margins and filled with calcite in their cores. Thin veins of pyrite cross-cut veins of prehnite, calcite, and Ca-zeolites. In general, the veins and vesicle fillings are consistent with sub-greenschist-, zeolite- or prehnite–actinolite-facies metamorphism.
Metamorphic conditions
Typical metamorphic mineral assemblages include albite, chlorite, prehnite, calcite, epidote, and titanite, ± Ca-zeolites (laumontite, thomsonite, heulandite), analcime, and actinolite. Minerals indicative of blueschist-facies metamorphism have not been observed. Metamorphic mineral assemblages typically found in the SFVC represent low temperatures and pressures of the zeolite to prehnite–actinolite subfacies. Analcime deposition and heulandite replacement of glass within the hyaloclastite layers probably occurred at temperatures that did not exceed 100°C (Shervais & Hanan, 1989). Zeolite minerals are stable only until 300°C and <0·1 GPa pressure. In general, rocks low in the section appear to have been metamorphosed at somewhat higher temperatures (300°C) than rocks near the top of the section (<100°C). Based on these conditions, it seems unlikely that the SFVC was ever emplaced within a subduction zone; in fact, the conditions of metamorphism observed are consistent with burial beneath sedimentary rocks of the Great Valley Series.
Tectonic setting of the SFVC
The SFVC contains three geochemically and petrologically distinct magma series, each of which provides a somewhat different view of the paleo-tectonic setting in which the complex formed. The discrimination plots of Leterrier et al. (1982) use pyroxene minor element chemistry to distinguish alkali basalts, non-orogenic basalts (MORB, OIB), and orogenic basalts (arc tholeiite, calc-alkaline; Fig. 16). These plots show that pyroxenes of the alkali suite are indeed alkaline in composition, whereas pyroxenes of the oceanic tholeiite and HALT suites are sub-alkaline (Fig. 16a). The sub-alkaline pyroxenes are further divided into non-orogenic (all of the oceanic tholeiite suite pyroxenes plus some HALT suite pyroxenes) and orogenic (the remaining HALT suite pyroxenes; Fig. 16b). Finally, the HALT suite pyroxenes straddle the dividing line between calc-alkaline and tholeiitic (Fig. 16c). The calc-alkaline character of the HALT suite is confirmed by MgO variation plots, which show that basalts of the HALT suite do not increase in FeO* and TiO2 with fractionation, consistent with calc-alkaline fractionation trends (Fig. 9).
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These three petrological series appear to represent two distinct tectonic settings. The Ti–Zr plot of Pearce & Cann (1973)
shows that the oceanic tholeiite and alkali suites are oceanic in character, but are more enriched in these elements than the limited data used to define this plot (Fig. 17a). The HALT suite has characteristics that overlap the arc tholeiite, calc-alkaline, and primitive MORB fields, with a significant gap between them and the other suites (Fig. 17a). The Ti–V plot of Shervais (1982) confirms these observations, and further shows that the Stonyford oceanic tholeiite and alkali suites are distinct: the Stonyford oceanic tholeiite suite has Ti/V ratios of 25–38, similar to MORB, whereas the alkali suite has Ti/V ratios of 45–54, similar to alkali basalts erupted in both continental and oceanic settings (Fig. 16b). In contrast, the HALT suite has Ti/V ratios that range from 15 to 30, overlapping the trends of both island arc volcanic suites and depleted MORB suites (Fig. 17b). Similar trends are seen in back-arc basin basalts, which vary from arc-like to oceanic in character.
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A plot of Ta vs Hf (Fig. 17c) shows that the oceanic tholeiite suite has Hf/Ta ratios similar to N-MORB (
15), but with much higher concentrations, consistent with small fractions of partial melting; this is consistent with the partial melting models presented earlier (Fig. 15). The alkali basalt series appears to lie on a mixing trend between E-MORB (Hf/Ta 4) and OIB (Hf/Ta 3); this trend suggests that a relatively enriched OIB-like source containing a recycled oceanic sediment component is needed to form these basalts (Fig. 17c). Finally, the HALT series basalts have Hf/Ta ratios that lie between N-MORB and E-MORB, but with concentrations that range from less than half to more than double these components (Fig. 17c). These trends are consistent with depletion of the HALT source mantle, followed by variable re-enrichment with a component similar in trace element character to the alkali basalt suite.
Relationship of the SFVC to the Franciscan assemblage in California
Shervais & Kimbrough (1985a, 1987) did not distinguish between the SFVC and the Snow Mountain Volcanic Complex of MacPherson (1983). However, as noted by MacPherson & Phipps (1985), these units are distinct and unrelated. The Snow Mountain Volcanic Complex contains incipient blueschist metamorphic phases and lies west of the serpentinite belt that marks the Coast Range fault (MacPherson, 1983; MacPherson & Phipps, 1985). In contrast, the SFVC contains unaltered volcanic glass, lacks incipient blueschist metamorphic phases, lies entirely within the Coast Range fault serpentinite belt, and is closely associated with (but is not directly overlain by) sediments of the Jurassic Knoxville formation (Shervais & Hanan, 1989; Zoglman, 1991). The SFVC also contains radiolarian chert intercalations with faunal assemblages that are similar to those found at other CRO locations, but distinct from faunal assemblages derived from Franciscan cherts (Shervais et al., 2005a). Thus, the Snow Mountain Volcanic Complex is clearly part of the Franciscan assemblage, whereas the SFVC is best regarded as part of the CRO (contrary to our former correlation of both units with the Franciscan assemblage, e.g. Shervais & Kimbrough, 1987; Shervais & Hanan, 1989; Shervais, 1990). This distinction is important because it has significant tectonic implications for the origin and evolution of the CRO.
Petrotectonic model for the SFVC and its relationship to Jurassic orogeny in California
The SFVC apparently overlies disrupted ophiolite similar to that found at Elder Creek and other Coast Range ophiolite localities, and recently published age dates show that it is slightly younger than the underlying ophiolite (Shervais et al., 2005a). Previously published geochemical data on these localities, along with data in unpublished theses, show that the CRO represents a supra-subduction zone ophiolite (Shervais & Kimbrough, 1985; Beaman, 1989; Shervais, 1990, 2001; Giaramita et al., 1998; Evarts et al., 1999; Snow, 2002; Shervais et al., 2004, 2005b). At Stonyford, remnants of ophiolitic rocks similar to those found at Elder Creek and other CRO localities form knockers in the serpentinite matrix mélange that underlies the volcanic complex. These knockers include wehrlite and pyroxenite cumulates, cumulate and isotropic gabbro, diorite and quartz diorite, and volcanic rocks. Their position beneath the SFVC suggests that they originally formed the substrate upon which the volcanic complex was erupted.
This conclusion is reinforced by radiolarian assemblages found in chert intercalations in the SFVC (Murchey, cited by Shervais et al., 2005a). Radiolaria found in the large lower chert lens above Stony Creek (near the base of the complex) represent the same assemblages as found in radiolarian chert that sits on top of CRO elsewhere. Radiolaria found in the upper chert lens above Dry Creek (near the top of the complex) correlate with radiolarian cherts that overlie the older cherts at other CRO localities (e.g. Hagstrum & Murchey, 1996). These observations require that the bulk of the SFVC was erupted onto its CRO substrate at the same time that ash-rich radiolarian chert was being deposited elsewhere, after formation of the main ophiolite sequence and prior to deposition of the Great Valley Series.
Where rocks with oceanic affinities are found at other CRO localities they always post-date formation of the main ophiolite (Shervais, 2001; Shervais et al., 2004). These include dikes with MORB-like geochemistry that intrude older SSZ ophiolite assemblages at Elder Creek, Black Mountain, Mount Diablo, Del Puerto, and Cuesta Ridge (Giaramita et al., 1998; Evarts et al., 1999; Shervais, 2001; Snow, 2002), inclusions of MORB glass in the Leona rhyolite (Shervais et al., 2004), and flows of MORB pillows that underlie chert and overlie SSZ volcanics at Cuesta Ridge (Snow, 2002). All of these occurrences suggest that after formation of the main ophiolite suites (arc tholeiite gabbros and volcanics, wehrlite–pyroxenite intrusions and boninitic lavas, and diorite–quartz diorite plutons and calc-alkaline volcanics) in an SSZ setting, the SSZ ophiolites interacted with a spreading ridge erupting N-MORB, E-MORB, and alkali basalts.
There are two situations where this is likely to occur: (1) where a back-arc basin spreading center propagates through an active arc setting, moving the active arc seaward and leaving a remnant arc behind (Karig, 1982); (2) collision of a spreading center with the subduction zone, where the spreading center is partly overridden by the actively extending forearc (Shervais, 2001; Shervais et al., 2004, 2005a). We prefer the latter explanation for several reasons. First, there are no arc rocks to the west that could represent either the active seaward arc or a remnant arc. This geometric problem persists regardless of which subduction polarity is assumed. Second, the rock series that form the main ophiolite assemblages of the CRO are consistent with formation in a rapidly extending forearc setting. These include the boninitic volcanic suites that are common at many CRO localities, and the equally common wehrlite–pyroxenite intrusive series. Harzburgites and dunites associated with the CRO at Stonyford contain depleted Cr-spinel (cr-numbers up to 74) that are characteristic of depleted forearcs parental to boninitic lavas (Shervais et al., 2005b).
Our preferred model is shown in Fig. 18: (a) formation of the CRO above a nascent subduction zone, as documented by previous studies; (b) collision of an active spreading center with this subduction zone, allowing influx of heterogeneous MORB-source mantle and melts derived from that mantle into the previously depleted supra-subduction mantle wedge; (c) finally, renewed subduction and shallow underthrusting to form the Franciscan complex and cause deformation associated with the Nevadan orogeny in the Sierra foothills (e.g. Shervais, 2001; Shervais et al., 2004). The geochemical data presented here and elsewhere, and the age data presented by Shervais et al. (2005a) contradict models that propose a far-traveled, exotic origin for the CRO as either a mid-oceanic ridge or back-arc basin assemblage (e.g. Hopson et al., 1981; Ingersoll, 2000; Pessagno et al., 2000).
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There are several lines of evidence supporting a model involving ridge collision at c. 160 Ma. Murchey & Blake (1993) found a reversal in faunal ages for accreted terranes in the Klamaths and the Sierras, such that residence time on the ocean floor (age of oceanic basement or oldest chert deposited on it versus time of arrival at the continental margin) of these accreted terranes became progressively shorter in the early Jurassic with a minimum at c. 160 Ma, followed by progressively longer residence times in the late Jurassic to early Cretaceous. They proposed that this minimum in residence time (<10 Myr) corresponds to a ridge collision event, followed by a long period of dominantly transcurrent relative motion (Murchey & Blake, 1993). Ward (1995)
showed that apparent polar wander paths for North America display an abrupt change at c. 160 Ma from longitudinal drift to rapid northward drift at the J2 cusp, a change that he attributed to a ridge collision and subsequent reordering of plate motions in response to the end of coupled motion between North America and the subducting plate. Finally, there is the change in kinematics of the Cordilleran margin shortly after the Nevadan phase of compression, from convergent in the middle Jurassic to sinistral transtension–transpression in the late Jurassic and early Cretaceous. |
ACKNOWLEDGEMENTS |
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This paper would not have been possible without the pioneering work and insights of Cliff Hopson, who introduced us (Shervais, Hanan) to the Coast Range ophiolite and who has provided the inspiration for our continued work there. This research was supported by NSF grants EAR8816398 and EAR9018721 (Shervais) and EAR9018275 (Kimbrough and Hanan). Geological mapping of the Stonyford ophiolite formed part of a Master's thesis by Marchell Zoglman Schuman (Zoglman, 1991) at the University of South Carolina. Instrumental neutron activation analyses were provided through the Oregon State University Reactor Sharing program (Robert J. Walker, analyst). Thoughtful reviews by Melanie Barnes, Dieter Mertz, and Reinhard Werner greatly improved the manuscript and are gratefully acknowledged.
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FOOTNOTES |
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Present address: 2623 Withington Peak Dr NE, Rio Rancho, NM 87144, USA. 
* Corresponding author. E-mail: shervais{at}cc.usu.edu
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REFERENCES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL RELATIONSPETROGRAPHY AND MINERALOGYWHOLE-ROCK CHEMISTRYISOTOPE GEOCHEMISTRYDISCUSSIONREFERENCES |
|---|
Arth, J. G. (1976). Behavior of trace elements during magmatic processes; a summary of theoretical models and their applications. Journal of Research of the US Geological Survey 4(1), 41–47.[Web of Science]
Beaman, B. J. (1989). Evolutional history of the plutonic section of the Elder Creek ophiolite: a remnant of the Coast Range ophiolite, California. Master's thesis, University of South Carolina, Columbia.
Ben Othman, D., White, W. M. & Patchett, P. J. (1989). The geochemistry of marine sediments, island arc magma genesis, and crust–mantle recycling. Earth and Planetary Science Letters 94, 1–21.[CrossRef][Web of Science]
Blake, M. C., Jr, Jayko, A. S., Jones, D. L. & Rogers, B. W. (1987). Unconformity between Coast Range Ophiolite and part of the lower Great Valley Sequence, South Fork of Elder Creek, Tehama County, California. In: Hill, M. L. (ed.) Cordilleran Section of the Geological Society of America, Centennial Field Guide. Boulder, CO: Geological Society of America, pp. 279–282.
Brown, R. D. (1964). Geologic map of the Stonyford quadrangle, Glenn, Colusa, and Lake counties, California. US Geological Survey Mineral Investigations Field Studies Map MF-279, scale 1:48 000.
Coish, R. A. & Taylor, L. A. (1979). The effects of cooling rate on texture pyroxene chemistry in DSDP Leg 34 basalt; a microprobe study. Earth and Planetary Science Letters 42(3), 389–398.[CrossRef][Web of Science]
Dennis, A. J. & Shervais, J. W. (1991). Style and significance of sheared serpentinite fault rocks, Stonyford, Northern California Coast Ranges. Geological Society of America, 1991 Annual Meeting 23(5), 236.
Dickinson, W. R., Hopson, C. A., Saleeby, J. B., Schweickert, R. A., Ingersoll, R. V., Pessagno, E. A., Jr, Mattinson, J. M., Luyendyk, B. P., Beebe, W., Hull, D. M., Munoz, I. M. & Blome, C. D. (1996). Alternate origins of the Coast Range Ophiolite, California; introduction and implications. GSA Today 6(2), 1–10.
Evarts, R. C., Coleman, R. G. & Schiffman, P. (1999). The Del Puerto ophiolite: petrology and tectonic setting. In: Wagner, D. L. & Graham, S. A. (eds) Geologic Field Trips in Northern California. California Division of Mines and Geology, Special Publication 119, 136–149.
Giaramita, M., MacPherson, G. J. & Phipps, S. P. (1998). Petrologically diverse basalts from a fossil oceanic forearc in California; the Llanada and Black Mountain remnants of the Coast Range Ophiolite. Geological Society of America Bulletin 110(5), 553–571.
Glen, R. A. (1990). Formation and thrusting in some Great Valley rocks near the Franciscan Complex, California, and implications for the tectonic wedging hypothesis. Tectonics 9(6), 1451–1477.[Web of Science]
Godfrey, N. J. & Klemperer, S. L. (1998). Ophiolitic basement to a forearc basin and implications for continental growth; the Coast Range/Great Valley Ophiolite, California. Tectonics 17(4), 558–570.[CrossRef][Web of Science]
Graham, D. W., Larsen, L. M., Hanan, B. B., Storey, M., Pedersen, A. K. & Lupton, J. E. (1998). Helium isotope composition of the early Iceland mantle plume inferred from the Tertiary picrites of West Greenland. Earth and Planetary Science Letters 160(3–4), 241–255.
Hagstrum, J. T. & Murchey, B. L. (1996). Paleomagnetism of Jurassic radiolarian chert above the Coast Range Ophiolite at Stanley Mountain, California, and implications for its paleogeographic origins. Geological Society of America Bulletin 108(6), 643–652.
Hanan, B. B. & Graham, D. W. (1996). Lead and helium isotope evidence from oceanic basalts for a common deep source of mantle plumes. Science 272, 991–995.[Abstract]
Hanan, B. B. & Schilling, J.-G. (1989). Easter Microplate evolution; Pb isotope evidence. Journal of Geophysical Research, B, Solid Earth and Planets 94(6), 7432–7448.[CrossRef]
Hemming, S. R. & McLennan, S. M. (2001). Pb isotope compositions of modern deep sea turbidites. Earth and Planetary Science Letters 184, 489–503.[CrossRef][Web of Science]
Hopson, C. A., Mattinson, J. M. & Pessagno, E. A. (1981). Coast Range ophiolite, western California. In: Ernst, W. G. (ed.) The Geotectonic Development of California: Rubey Volume 1. Washington, DC: American Geophysical Union, pp. 418–510.
Huot, F. & Maury, R. C. (2002). The Round Mountain serpentinite melange, northern Coast Ranges of California: an association of backarc and arc-related tectonic units. Geological Society of America Bulletin 114(1), 109–123.
Ingersoll, R. V. (2000). Models for origin and emplacement of Jurassic ophiolites of northern California. In: Dilek, Y., Moores, E. M., Elthon, D. & Nicolas, A. (eds) Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America, Special Papers 349, 395–402.
Janney, P. E. & Castillo, P. R. (1997). Geochemistry of Mesozoic Pacific mid-ocean ridge basalt; constraints on melt generation and the evolution of the Pacific upper mantle. Journal of Geophysical Research, B, Solid Earth and Planets 102(3), 5207–5229.
Jayko, A. S., Black, M. C. & Harms, T. (1987). Attenuation of the Coast Range ophiolite by extensional faulting and the nature of the Coast Range thrust, California. Tectonics 6, 475–488.[Web of Science]
Karig, D. E. (1982). Initiation of subduction zones; implications for arc evolution and ophiolite development. In: Leggett, J. K. (ed.) Trench–Forearc Geology: Sedimentation and Tectonics on Modern and Ancient Active Plate Margins. Geological Society, London, Special Publications 10, 563–576.
Laul, J. C. (1986). Chemistry of the Apollo 12 highland component. Proceedings of the 16th Lunar and Planetary Science Conference. Journal of Geophysical Research, 91(B4), D251–D261.
Le Maitre, R. W. (1981). GENMIX—a generalized petrological mixing model program. Computers and Geosciences 7, 229–247.[CrossRef]
Leterrier, J., Maury, R. C., Thonon, P., Girard, D. & Marchal, M. (1982). Clinopyroxene composition as a method of identification of the magmatic affinities of paleo-volcanic series. Earth and Planetary Science Letters 59(1), 139–154.[CrossRef][Web of Science]
MacPherson, G. J. (1983). The Snow Mountain volcanic complex; an on-land seamount in the Franciscan terrain, California. Journal of Geology 91(1), 73–92.[Web of Science]
MacPherson, G. J. & Phipps, S. P. (1985). Geochemical evidence for the tectonic setting of the Coast Range ophiolite; a composite island arc–oceanic crust terrane in western California; discussion. Geology, Boulder 13(11), 827–828.[CrossRef]
McKenzie, D. & O'Nions, R. K. (1991). Partial melt distributions from inversion of rare earth element concentrations. Journal of Petrology 32, 1021–1091.
McKenzie, D. & O'Nions, R. K. (1995). The source regions of ocean island basalts. Journal of Petrology 36, 133–159.
McLennan, S. M., Taylor, S. R., McCulloch, M. T. & Maynard, J. B. (1990). Geochemical and Nd–Sr isotopic composition of deep-sea turbidites: crustal evolution and plate tectonic associations. Geochimica et Cosmochimica Acta 54, 2015–2050.[CrossRef][Web of Science]
Mertz, D. F., Weinrich, A. J., Sharp, W. D. & Renne, P. R. (2001). Alkaline intrusions in a near-trench setting, Franciscan Complex, California; constraints from geochemistry, petrology, and 40Ar/39Ar chronology. American Journal of Science 301(10), 877–911.
Metcalf, R. V., Wallin, E. T., Willse, K. R. & Muller, E. R. (2000). Geology and geochemistry of the ophiolitic Trinity Terrane, California; evidence of middle Paleozoic depleted supra-subduction zone magmatism in a proto-arc setting. In: Dilek, Y., Moores, E. M., Elthon, D. & Nicolas, A. (eds) Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America, Special Papers 349, 403–418.
Moores, E. M. (1982). Origin and emplacement of ophiolites. Reviews of Geophysics and Space Physics 20(4), 735–760.[Web of Science]
Murchey, B. L. & Blake, M. C., Jr (1993). Evidence for subduction of a major ocean plate along the California margin during the Middle to Early Late Jurassic. In: Dunne, G. C. & McDougall, K. (eds) Mesozoic Paleogeography of the Western United States II, Vol. 71. Los Angeles, CA:Society of Economic Paleontologists and Mineralogists, Pacific Section, pp. 1–17.
Palfy, J., Smith, P. L. & Mortensen, J. K. (2000). A U–Pb and 40Ar/39Ar time scale for the Jurassic. Canadian Journal of Earth Sciences 37(6), 923–944.[CrossRef][Web of Science]
Pearce, J. A. & Cann, J. R. (1973). Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth and Planetary Science Letters 19(2), 290–300.[CrossRef][Web of Science]
Pessagno, E. A., Jr, Hull, D. M. & Hopson, C. A. (2000). Tectonostratigraphic significance of sedimentary strata occurring within and above the Coast Range Ophiolite (California Coast Ranges) and the Josephine Ophiolite (Klamath Mountains), northwestern California. In: Dilek, Y., Moores, E. M., Elthon, D. & Nicolas, A. (eds) Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America, Special Papers 349, 383–394.
Pouchou, J.-L. & Pichoir, F. (1991). Quantitative analysis of homogeneous or stratified microvolumes applying the model ‘PAP’. In: Heinrich, K. F. J. & Newbury, D. E. (eds) Electron Probe Quantitation. New York: Plenum, pp. 31–76.
Saleeby, J. B. (1992). Petrotectonic and paleogeographic settings of U.S. Cordilleran ophiolites. In: Burchfiel, C., Lipman, P. W. & Zoback, M. L. (eds) The Cordilleran Orogen; Conterminous U.S. Boulder, CO: Geological Society of America, pp. 653–682.
Schweitzer, E. L., Papike, J. J. & Bence, A. E. (1979). Statistical analysis of clinopyroxenes from deep-sea basalts. American Mineralogist 64(5–6), 501–513.[Abstract]
Shervais, J. W. (1982). Ti–V plots and the petrogenesis of modern and ophiolitic lavas. Earth and Planetary Science Letters 59(1), 101–118.[CrossRef][Web of Science]
Shervais, J. W. (1990). Island arc and ocean crust ophiolites: contrasts in the petrology, geochemistry, and tectonic style of ophiolite assemblages in the California Coast Ranges. In: Malpas, J., Moores, E. M., Panayiotou, A. & Xenophontos, C. (eds) Ophiolites: Oceanic Crustal Analogues, Proceedings of the Symposium Troodos 1987, Nicosia: Geological Survey Department, pp. 507–520.
Shervais, J. W. (2001). Birth, death, and resurrection: the life cycle of supra-subduction zone ophiolites. Geochemistry, Geophysics, Geosystems 2, paper number 2000GC000080.
Shervais, J. W. & Hanan, B. B. (1989). Jurassic volcanic glass from the Stonyford volcanic complex, Franciscan assemblage, Northern California Coast Ranges. Geology, Boulder 17(6), 510–514.
Shervais, J. W. & Kimbrough, D. L. (1985a). Geochemical evidence for the tectonic setting of the Coast Range ophiolite; a composite island arc–oceanic crust terrane in western California. Geology, Boulder 13(1), 35–38.
Shervais, J. W. & Kimbrough, D. L. (1985b). Geochemical evidence for the tectonic setting of the Coast Range ophiolite; a composite island arc–oceanic crust terrane in western California; reply. Geology, Boulder 13(11), 828–829.[CrossRef]
Shervais, J. W. & Kimbrough, D. L. (1987). Alkaline and transitional subalkaline metabasalts in the Franciscan Complex melange, California. In: Morris, E. M. & Pasteris, J. D. (eds) Mantle Metasomatism and Alkaline Magmatism. Geological Society of America, Special Papers 215, 165–182.
Shervais, J. W., Vetter, S. K. & Lindstrom, M. M. (1990). Chemical differences between small subsamples of Apollo 15 olivine-normative basalts. In: Sharpton, V. L. & Ryder, G. (eds) Proceedings of the 20th Lunar and Planetary Science Conference. Houston, TX: Lunar and Planetary Institute, pp. 109–126.
Shervais, J. W., Murchey, B., Kimbrough, D. L., Hanan, B. B., Renne, P. R., Snow, C. A., Schuman, M. Z. & Beaman, J. (2004). Multi-stage origin of the Coast Range Ophiolite, California: implications for the life cycle of supra-subduction zone ophiolites. International Geology Review 46, 289–315.[Web of Science]
Shervais, J. W., Murchey, B., Kimbrough, D. L., Renne, P. & Hanan, B. B. (2005a). Radioisotopic and biostratigraphic age relations in the Coast Range Ophiolite, Northern California: implications for the tectonic evolution of the Western Cordillera. Geological Society of America Bulletin 117(5/6), 633–653.
Shervais, J. W., Kolesar, P. & Andreasen, K. (2005b). Field and chemical study of serpentinization—Stonyford, California: chemical fluxes and mass balance. International Geology Review 47, 1–23.[Web of Science]
Sigurdsson, H. & Schilling, J.-G. (1976). Spinels in mid-Atlantic Ridge basalts: chemistry and occurrence. Earth and Planetary Science Letters 29, 7–20.[CrossRef][Web of Science]
Simpson-Seymore, A. (1999). Tectonic implications of the Stony Creek basaltic sandstone, Late Jurassic Great Valley Series, Stonyford, California. M.Sc. thesis, University of South Carolina, Columbia, 122 pp.
Snow, C. A. (2002). Geology of the Cuesta Ridge ophiolite remnant near San Luis Obispo, California: evidence for the tectonic setting and origin of the Coast Range Ophiolite. M.S. thesis, Utah State University, Logan.
Stakes, D. S., Shervais, J. W. & Hopson, C. A. (1984). The volcanic–tectonic cycle of the FAMOUS and AMAR valleys, Mid-Atlantic Ridge, 36°47'N; evidence from basalt glass and phenocryst compositional variations for a steady state magma chamber beneath the valley midsections, AMAR 3. Journal of Geophysical Research, B, Solid Earth and Planets 89(8), 6995–7028.[CrossRef]
Stern, R. J. & Bloomer, S. H. (1992). Subduction zone infancy; examples from the Eocene Izu–Bonin–Mariana and Jurassic California arcs. Geological Society of America Bulletin 104(12), 1621–1636.
Sun, S.-S. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications 42, 313–345.
Sun, S.-S., Tatsumoto, M. & Schilling, J.-G. (1975). Mantle plume mixing along the Reykjanes ridge axis: lead isotope evidence. Science 190, 143–147.[Web of Science]
Thompson, R. N., Morrison, M. A., Dickin, A. P. & Hendry, G. L. (1983). Continental flood basalts ... arachnids rule OK? In: Hawkesworth, C. J. & Norry, M. J. (eds) Continental Basalts and Mantle Xenoliths. Nantwich: Shiva, pp. 158–185.
Tilling, R. I., Wright, T. L. & Millard, H. T., Jr (1987). Trace element chemistry of Kilauea and Mauna Loa lava in space and time: a reconnaissance. US Geological Survey, Professional Papers 1350, 641–689.
Todt, W., Cliff, R. A., Hanser, A. & Hofmann, A. W. (1984). 202Pb + 205Pb double spike for lead isotopic analyses. Terra Cognita 4, 209.
Ward, P. L. (1995). Subduction cycles under western North America during the Mesozoic and Cenozoic eras. In: Miller, D. M. & Busby, C. (eds) Jurassic Magmatism and Tectonics of the North American Cordillera. Geological Society of America, Special Papers 299, 1–45.
Wentworth, C. M., Blake, M. C., Jr, Jones, D. L., Walter, A. W. & Zoback, M. D. (1984). Tectonic wedging associated with emplacement of the Franciscan assemblage, California Coast Ranges. In: Blake, C., Jr (ed.) Franciscan Geology of Northern California, Vol. 43. Los Angeles, CA: Society of Economic Paleontologists and Mineralogists, Pacific Section, pp. 163–173.
White, W. M. (1993). 238U/204Pb in MORB and open system evolution of the depleted mantle. Earth and Planetary Science Letters 115, 211–226.[CrossRef][Web of Science]
Wright, T. L. (1971). Chemistry of Kilauea and Mauna Loa lava in space and time. US Geological Survey, Professional Papers 735, 40 pp.
Zoglman, M. M. (1991). The Stonyford Volcanic Complex: petrology and structure of a Jurassic seamount in the Northern California Coast Ranges. M.S. thesis, University of South Carolina, Columbia.
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