Journal of Petrology Advance Access published online on March 14, 2007
Journal of Petrology, doi:10.1093/petrology/egm003
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||
Diverse Origins of Xenoliths from Seamounts at the Continental Margin, Offshore Central California
Monterey Bay Aquarium Research Institute, 7700 Sandholdt Road, Moss Landing, CA 95039-9644, USA
Received August 10, 2005; Revised typescript accepted January 31, 2007
 |
ABSTRACT |
|---|
 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
A diverse assemblage of small mafic and ultramafic xenoliths occurs in alkalic lava from Davidson and Pioneer seamounts located at the continental margin of central California. Based on mineral compositions and textures, they form three groups: (1) mantle xenoliths of lherzolite, pyroxenite, and dunite with olivine of >Fo90; (2) ocean crust xenoliths of dunite with olivine <Fo90, troctolite, pyroxene-gabbro, and anorthosite with low-K2O plagioclase; (3) cumulates of seamount magmas of alkalic gabbro with primary amphibole and biotite and anorthosites with high-K2O plagioclase. The alkalic cumulates are genetically related to, but more evolved than, their host lavas and probably crystallized at the margins of magma reservoirs. Modeling and comparison with experimentally derived phases suggest an origin at moderate pressures (
0·5–0·9 GPa). The high volatile contents of the alkalic host lavas may have pressurized the magma chambers and helped to propel the xenolith-bearing lavas directly from deep storage at the base of the lithosphere to the eruption site on the ocean floor, entraining fragments of the upper mantle and ocean crust cumulates from the underlying abandoned spreading center. KEY WORDS: basaltic magmatism; continental margin seamounts; geothermobarometry; mineral chemistry; xenoliths
|
INTRODUCTION |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Xenoliths are samples of the mantle lithosphere underlying a volcano and/or the crust the host magma traversed. Their compositions provide information about the temperatures and pressures at which they originated or last equilibrated. They might record metasomatic processes that modify the lower lithosphere during rock–melt interactions. Mantle xenoliths have been described from various tectonic settings, including continental rifts (e.g. Frey & Prinz, 1978
; Kempton, 1987; McGuire, 1988), island arcs (e.g. Takahashi, 1980), and ocean islands such as Hawaii (e.g. Sen & Presnall, 1986; Sen, 1988; Sen et al., 2005), the Canary Islands (e.g. Neumann, 1991; Neumann et al., 2000, and references therein), and the Society Islands (Qi et al., 1994). Some ocean island lavas containing mantle xenoliths also include fragments of old ocean crust (e.g. Clague & Chen, 1986; Fodor & Vandermeyden, 1988; Schmincke et al., 1998; Neumann et al., 2000). Other xenoliths are cumulates of ocean island (e.g. Sen & Presnall, 1986; Clague, 1987; Bohrson & Clague, 1988; Fodor & Moore, 1994; Fodor & Galar, 1997) or mid-ocean ridge magma chambers (e.g. Hekinian et al., 1985; Dixon et al., 1986; Davis & Clague, 1990). If the xenoliths ascend rapidly, there might be minimal interaction with their host magma. In contrast, if they are in prolonged contact with the melt, the xenoliths might be mineralogically and chemically modified.
This study describes the petrography and mineral and host lava compositions of a diverse suite of xenoliths from Davidson and Pioneer seamounts, offshore central California. Unlike most intra-plate ocean island volcanoes, the seamounts are built on top of spreading center segments that were abandoned at the continental margin when the tectonic regime changed from subduction to a transform margin. The host lavas erupted millions of years after mid-ocean ridge basalt (MORB) volcanism ended (Davis et al., 1995, 2002). The xenoliths provide a window into the upper mantle and lower crust in this unusual environment. We use the mineral chemistry to identify and distinguish mantle and ocean crust cumulates from xenoliths related to the alkalic volcanism that built the seamounts. We estimate the depth of origin based on temperatures and pressures recorded by mineral equilibria in the xenoliths and draw inferences concerning magma generation and transport processes.
|
GEOLOGICAL SETTING |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Davidson and Pioneer are two seamounts of a group of four located at the continental margin, offshore central California (Fig. 1). Morphologically similar seamounts are more abundant offshore southern and Baja California. Unlike typical ocean island volcanoes or near-ridge seamounts, all of the seamounts are complex NE–SW-trending ridges that reflect the ridge-parallel structure of the underlying oceanic crust (Davis et al., 2002
). Davidson Seamount is built on a fossil spreading center based on symmetric magnetic anomalies of Chron 6, indicating an ocean crust age of about 20 Ma (Lonsdale, 1991). Mapped Chron 6C magnetic anomalies near Pioneer Seamount are not symmetrical about the seamount but suggest an age of about 24 Ma for the underlying ocean crust. Published 39Ar/40Ar laser fusion ages indicate volcanism at about 12 Ma on Davidson and at 11 Ma on Pioneer (Davis et al., 2002), younger by 8–13 Myr than the underlying oceanic crust. New Ar–Ar incremental heating results for some Davidson samples expand the age of volcanism at Davidson from 17 to 10 Ma (D. A. Clague, unpublished data), indicating that episodes of volcanism occurred on 3 to 10 Myr old ocean crust. Such prolonged volcanic activity to form the seamounts suggests very low magma supply rates and long hiatuses between eruptions, as suggested based solely on seamount morphology by Davis et al. (2002).
|
Some whole-rock and glass chemistry data from Davidson and Pioneer seamounts were given by Davis et al. (2002
), who presented petrography, Ar–Ar ages, and trace element and isotope compositions for lavas from the four seamounts offshore central California and for one located farther south. Volcanic rocks are predominantly alkalic basalt, hawaiite, and mugearite, but also include some tholeiitic basalt and rare trachyte. Radiogenic isotopes indicate a variably enriched MORB source (Davis et al., 2002; P. Castillo, personal communication). |
SAMPLING AND ANALYTICAL METHODS |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
The xenoliths occur in volcanic rocks that were collected by dredging on several cruises of the US Geological Survey (USGS) in 1976, 1978, and 1979 and on dives of the remotely operated vehicle (ROV) Tiburon on three cruises of the Monterey Bay Aquarium Research Institute's R.V. Western Flyer in 2000 and 2002. The xenoliths studied were selected to include the widest variety of minerals and textures, but they represent only a small fraction of the inclusions present in the lavas. Whole-rock lava samples were analyzed by X-ray fluorescence (XRF) at the GeoAnalytical Laboratory of Washington State University, and the standards used, precision and accuracy are available at their web site (http://www.wsu.edu/~geology/geolab/note.html). Minerals of xenoliths and glass of pillow rinds and of volcanic breccias were analyzed with a JEOL 8900 Superprobe at the USGS, Menlo Park using natural and synthetic glass and mineral standards (Davis et al., 1994
). Glass and plagioclase were analyzed with a defocused beam (10 µm) and 20 nA and 15 nA specimen current, respectively. A focused beam and 25 nA were used for pyroxene, amphibole, biotite, and apatite; for olivine and oxides the current was increased to 30 nA. Back-scattered electron images of textural features and compositional zoning were also determined with the same microprobe rasterizing with a focused beam and 40 nA current over a variable-sized area, according to the area of interest. The complete analytical dataset is available at http://petrology.oxfordjournals.org (Electronic Appendices 1–3). |
HOST LAVA |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
The host lavas containing the xenoliths are alkalic basalt, hawaiite and mugearite (Table 1, Fig. 2). The rare tholeiitic basalt and trachyte do not contain xenoliths. One calc-alkaline andesite that was also recovered is not included here because it is inferred to be an erratic. Exotic rocks, including granitic, sedimentary, and metamorphic rocks, occur on all of these seamounts (Davis et al., 2002
; Paduan et al., 2004). Several of the xenoliths occur in fresh or altered glass of volcanic breccia. The xenolith-bearing, alkalic lavas are moderately to highly vesicular and are typically porphyritic with variable proportions of clinopyroxene, plagioclase, ±olivine. Broken fragments of large plagioclase crystals in a number of samples may be pieces of xenoliths. The rims of these crystals are more calcic than the cores and overlap with the compositions of microphenocrysts and microlites (>An50, K2O 0·3–0·7 wt%). The clinopyroxene is pinkish brown, typically complexly zoned, with high TiO2 (to 7 wt%) and Al2O3 (to 13 wt%) diagnostic of alkalic basalts (LeBas, 1962). The olivine is often replaced by clays and iron oxides but when unaltered is Fo78–87 with 0·2–0·3 wt% CaO and 0·03–0·25 wt% NiO.
|
|
Xenoliths are found in different flows at numerous locations on the seamounts (Fig. 1). A wide range of xenoliths, including dunite, lherzolite, pyroxenite, amphibole-gabbro, anorthosite, and amphibole and titanomagnetite megacrysts has been found; these are especially abundant in one mugearite sample (L2-79NC-D1-R35) recovered in a dredge from the northern flank of Davidson Seamount.
Representative major element compositions of 54 whole-rock samples and 93 glasses are listed in Table 1 and shown in Fig. 2. Whole-rock compositions range from tholeiitic and mildly alkalic basalt to trachyte but hawaiite compositions are most abundant (Fig. 2). Glass rims of lava are typically more evolved than their corresponding whole-rock compositions, but trends parallel those of the whole-rock samples. The dense, aphyric trachyte has no glass rind. Similar to other ocean island suites (e.g. fig. 1 of Nevkasil et al., 2004, and references therein), the greatest variability, especially in alkalis, occurs over a narrow range of silica contents ( 48–50 wt%, Fig. 2). More evolved compositions show better developed trends, but they also represent a limited number of samples. MgO, CaO, FeO, and TiO2 decrease with increasing SiO2 whereas Na2O and K2O, and to a lesser extent Al2O3, increase. P2O5 has the greatest scatter (figures not shown).
|
DESCRIPTION OF XENOLITHS AND MEGACRYSTS |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Petrographic descriptions are summarized in Table 2, images of representative thin sections are shown in Fig. 3, and back-scattered electron (BSE) images of selected areas are shown in Fig. 4. Xenoliths include dunite, lherzolite, pyroxenite, troctolite, anorthosite, and gabbro. The gabbro xenoliths can be divided into three groups based on the presence or absence of amphibole and whether the amphibole is primary or secondary. Megacrysts are plagioclase, amphibole, clinopyroxene, and titanomagnetite (Table 2, Fig. 3). Except for two 10 cm gabbro samples from Davidson Seamount, xenoliths are small (<0·5 to 5 cm) inclusions in crystalline or glassy basaltic lava or volcaniclastic breccia. Megacrysts are single large crystals ranging in size from
4 to 9 mm that were identified as xenocrysts based primarily on mineral compositions, discussed below.
|
|
|
Dunite
The five analyzed dunite xenoliths (Table 2) are small (
5 to 8 mm) and angular with some planar surfaces and/or rounded corners (Fig. 3a). All have fractures and joints in two or more directions; some are severely sheared. Brown iron oxide alteration lines many of the fractures and replaces some olivine, which is commonly strained. Textures are predominantly fine- to medium-grained, allotriomorphic–granular (Pike & Schwarzman, 1977), except for sample D3 (Fig. 4a), which has a narrow layer of a mosaic of small, anhedral olivine crystals at the margin. Other than this margin, contacts with the host lavas are typically sharp and mostly without reaction rims. Three samples contain subhedral to rounded Cr-rich spinel crystals to 0·5 mm in size. Only two dunites (D3, D5) have olivine >Fo90 (Table 3, Fig. 5, and Electronic Appendix 1). Olivine compositions in the other three samples have Fo87–88 and the small crystals forming the mosaic at the margin of D3 is Fo86±1. Except for these small crystals with 0·26 wt% CaO, all are typically low in CaO (0·01–0·05 wt%) relative to olivine in the host lava (Fig. 5a). NiO content is >0·30 wt% for samples with >Fo90 (Table 3, Fig. 5b) but lower Fo olivine has correspondingly lower NiO (0·20–0·30 wt%). The spinel, occurring only in the three samples with lower Fo olivine, has TiO2 < 0·4 wt% and Al2O3 contents of about 30 wt% or less (Table 4).
|
|
|
Lherzolite and pyroxenite
Four lherzolite xenoliths range from triangular to blocky in shape (Table 2, Fig. 3b). There are few if any reaction rims, although abundant blebs and stringers of lava have penetrated into some of the xenoliths along fractures (Fig. 4b). Fractures or joints in at least two directions are typically present. Only sample L1 (Fig. 3b) is deformed, although it does not have well-developed foliation. This sample is the only one to which the textural term porphyroclastic (Pike & Schwarzman, 1977
) is applicable. Olivine crystals are strained and may be partially or completely replaced by iddingsite and/or iron oxide. Orthopyroxene crystals typically have undeformed clinopyroxene exsolution lamellae (Fig. 4b). Spinel is present in only two samples (L1, L2) in the form of minute, anhedral crystals along pyroxene crystal boundaries.
The olivine compositions are comparable with those in the dunites (Table 3, Fig. 5, and Electronic Appendix 1). One sample has Fo87–88 and the other three have >Fo90. CaO ranges from 0·01 to 0·06 wt% and NiO from 0·30 to 0·45 wt%. Clinopyroxene is calcic Cr-diposide, and orthopyroxene is enstatite (Fig. 6a). Both pyroxenes have high Mg-numbers ranging from 87 to 94 (Fig. 6b and c). TiO2 in clinopyroxene is low (0·00–0·4 wt%) and Cr2O3 is high, up to 2·9 wt% (Table 5, Electronic Appendix 2), relative to that of the host lavas. Al2O3 in clinopyroxene is also typically low, ranging from 0·8 to 3 wt%. Only small subhedral clinopyroxene neoblasts along the margins and in the fractures of the deformed lherzolite (L1) are more aluminous (4–7 wt%). Despite the somewhat higher Al2O3, they are uniformly low in TiO2 (0·03 wt%, Fig. 6c) and have some of the highest Cr2O3 contents (to 2·6 wt%). Spinel in sample L3 is more aluminous than in the dunites (37 wt%) but lower in TiO2 (0·07 wt%). Rare spinel in the deformed lherzolite (L1) is the most Cr2O3-rich (47·5 wt%, Table 4, Fig. 7).
|
|
|
The two pyroxenites are small (<6 mm), angular slivers of mostly orthopyroxene with clinopyroxene occurring as small crystals and as undeformed exsolution lamellae. One sample (P1) has euhedral spinel inclusions up to 1 mm in size (Fig. 3c). The pyroxenes have Mg-numbers ranging from 86 to 94 (Table 5, Electronic Appendix 2). The clinopyroxene is low in TiO2 (to 0·17 wt%) and Al2O3 (<4 wt%, Fig. 6c). Spinel compositions are similar to those in dunite and lherzolite, having low TiO2 (<0·10 wt%) and Al2O3 contents of
32 wt% (Table 4).
Troctolite
One xenolith consists of plagioclase and olivine, which encloses small (<0·1 mm), subhedral spinel crystals. The angular, 1·5 cm fragment has an equigranular texture (Fig. 3d). There is no reaction rim and virtually no alteration of either olivine or plagioclase crystals. The plagioclase is labradorite (An58–65) with a low K2O (0·10 wt%) content (Table 6, Fig. 8). The olivine is Fo85 with a high NiO (to 0·3 wt%) and low CaO (<0·07 wt%) content. The small Cr-rich spinel crystals are highly aluminous (37·5–43·5 wt%) and have low TiO2 (0·03 wt%) contents.
|
|
Gabbro without amphibole
Three gabbro xenoliths do not contain amphibole. One of these (Gn1) is composed of plagioclase and clinopyroxene with traces of olivine replaced by iddingsite and Fe-oxide. It is medium to coarse grained (
5 mm). Contact with the host lava is sharp and without reaction rims. Only one fracture extends through this xenolith and into the host lava (Fig. 3e). As no lava has penetrated into the fracture, fracturing must have occurred post-emplacement, possibly during sample collection or preparation. The plagioclase is low-K2O (0·20 wt%) labradorite, and the clinopyroxene is low in TiO2 (0·5–1·5 wt%) and Al2O3 (3·7–6 wt%) relative to pyroxene in the host lava (Fig. 6b). The Mg-numbers of clinopyroxene in Gn1 range from 77 to 80 (Table 5, Electronic Appendix 2).
The two other gabbro samples (Gn2, Gn3), the largest of all the xenoliths (10 cm), are composed of clinopyroxene, plagioclase, and minor olivine. They have basically the same lithology except that one has a larger proportion of pyroxene relative to plagioclase than the other. Both samples are almost disaggregated by their host lava (Fig. 3f). The plagioclase is highly anhedral with embayed margins and zones of sieve-texture, with some crystals having a narrow rim of a more calcic overgrowth. Compositions are higher in K2O than for Gn1, ranging from An54 and 0·4 wt% K2O (Table 6) for cores to An70 and 0·3 wt% K2O (Fig. 8) for rims. The clinopyroxene crystals are also anhedral with severely embayed margins and devitrified glass inclusions. Most crystals are optically zoned, and Mg numbers range from 71 to 76·7. The TiO2 (1· 4–4·2 wt%) and Al2O3 (6·2–11·2 wt%) contents of clinopyroxene are significantly higher than those of Gn1 (Fig. 6c), indicating that they crystallized from an alkalic melt. The unaltered olivine is Fo76–80 and NiO ranges from 0·02 to 0·24 wt% and CaO from 0·14 to 0·25 wt% (Table 3, Fig. 5). The compositions of olivine cores in these two xenoliths are lower in Fo and CaO, but similar in NiO, to those of the rims, which overlap with the olivine composition in the surrounding lava (Fig. 5a).
Gabbro with secondary amphibole
Three gabbro xenoliths contain dark brown, dusty-looking amphibole that we interpret to be secondary because it occurs as discontinuous, anhedral inclusions within clinopyroxene, or as a replacement along crystal margins and in fractures (Fig. 3g). All three samples are medium- to coarse-grained, allotriomorphic–granular in texture and have extensively reacted with the host melt. Two (Gs1, Gs2) are composed primarily of plagioclase, clinopyroxene, and minor olivine, replaced by iddingsite. One sample (Gs2) has pyroxene with crisscrossing Fe–Ti oxide lamellae (Fig. 4c) that may have replaced orthopyroxene. A third sample (Gs3) has relict orthopyroxene lamellae that are too narrow to analyze. The feldspar in Gs1 and Gs2 is labradorite (An52–65) with a low K2O content (0·14–0·22 wt%), whereas Gs3 has some core compositions of labradorite of An>50, K2O 0·3 wt% but rims are more sodic and have K2O up to 1 wt% (Fig. 8). Rim compositions of plagioclase in the xenoliths are comparable with those in the host lava and probably reacted with the host melt. The unaltered clinopyroxene cores of Gs1 and Gs2 have Mg-numbers ranging from 72 to 91. TiO2 and Al2O3 range from 0·6 to 1·6 wt% and 2·8 to 5·5 wt%, respectively. The clinopyroxene in Gs3 is too altered to analyze. The amphiboles in all three samples have SiO2 of 41 ± 1 wt%, with TiO2 ranging from 3·7 to 5·4 wt% and K2O from 0·9 to 1·34 wt% (Table 7). The lower K2O and TiO2 compositions are limited to round or anhedral inclusions in the clinopyroxene. Only these three gabbro xenoliths show extensive mineral replacement indicative of modal metasomatism (e.g. Kempton, 1987).
|
Gabbro with primary amphibole
Eight gabbro xenoliths consist of medium- to coarse-grained, hypidiomorphic–granular aggregates of predominantly plagioclase and brown, strongly pleochroic amphibole; only one of these (Gp5) also contains clinopyroxene. We interpret the amphibole to be magmatic in origin because it occurs as optically homogeneous, large crystals (to 0·9 cm), bounded in part by crystal faces, and because it poikilitically encloses plagioclase. Four samples are mostly plagioclase with minor amphibole inclusions or anhedral amphibole attached at the margins, whereas two others (Gp3, Gp6) consists mostly of amphibole poikilitically enclosing small, euhedral plagioclase crystals (Fig. 3h). The amphibole of Gp5 forms coronas around the clinopyroxene. Traces of anhedral biotite are present along the margins of amphibole crystals in three samples and euhedral apatite and Fe–Ti oxide are included in plagioclase of several samples (Table 2). Rare iron sulfide (pyrrhotite) inclusions occur in some amphiboles. One of the gabbro samples (Gp1) is almost disaggregated by the host lava (Fig. 3i). The amphibole, spanning a narrow compositional range (Table 7), is kaersutite [classification of Leake (1978
)] with SiO2 ranging from 38 to 39·7 wt%, TiO2 from 4·5% to 8 wt%, and K2O from 1·1 to 1·4 wt% (Fig. 9). The biotite has 8–9 wt% K2O and 4–8% wt TiO2. The clinopyroxene in sample Gp5 is moderately high in TiO2 (1·9 wt%) and Al2O3 (7·5 wt%), comparable with some in the host lavas. The feldspar compositions are predominantly andesine but range from labradorite to oligoclase with K2O from 0·21 to 1·0 wt%. Reverse zoning is ubiquitous, and the edges of feldspar crystals in contact with the host lava have an overgrowth of labradorite, comparable with the compositions of small crystals in the host lava (Fig. 8). Oxide present is mostly titanomagnetite, but one sample has both titanomagnetite and ilmenite (Gp8), and a third has only ilmenite (Gp2, Table 8). Euhedral apatite crystals are enclosed in feldspar and have F contents to 1· 4 wt% and Cl to 0·5 wt% (Table 9).
|
|
|
Anorthosite
Eleven xenoliths consist primarily of feldspar. Inclusions of euhedral apatite and subhedral titanomagnetite are present in or attached to the feldspar of six samples. In one sample (A2), feldspar encloses a large (0·8 mm) zircon crystal (Table 2). The anorthosite xenoliths range in size from about 1 to 4 cm and occur in two basic shapes: nearly elliptical with rounded corners (Fig. 3j) or elongated–tabular with deeply embayed margins that are aligned along cleavage planes (Fig. 3k). Fractures in two or more directions are often present. Devitrified glass inclusions in feldspar are highly abundant in some samples. Reaction rims are virtually absent in some samples, whereas others have embayed and sieve-textured margins and alteration along fractures and cleavage planes.
Compositionally, the feldspar is predominantly andesine but ranges from An65 to An13 with correspondingly increasing K2O (0·30–1·85 wt%, Fig. 8). Reaction rims are more pronounced in samples with more sodic feldspar and are most severe for the oligoclase (An13) of the zircon-bearing xenolith (A2). All analyzed crystals are reversely zoned, including labradorite crystals that show no evidence of resorption. As observed for the primary amphibole-gabbros, titanomagnetite and apatite inclusions occur only in the more sodic plagioclase and have compositions (Tables 8 and 9) comparable with those in the gabbros with primary amphibole. Coexisting ilmenite and titanomagnetite were found in one sample (A9).
Megacrysts
Large, single crystals of amphibole, titanomagnetite, feldspar, and clinopyroxene present in some lava samples appear to be xenocrysts, based on compositions. Distinctive, large (to 0·9 cm) rounded amphibole crystals (Fig. 3l) are present in hawaiite and mugearite lava samples (Table 2) that also contain amphibole-gabbro xenoliths. Similar in size and pleochroism, and with comparable kaersutite compositions (TiO2 5·5 to 7 wt% and K2O 1·2 wt%, Fig. 9), they appear to be disaggregated xenoliths. Their shape is largely ovoid and some have embayed margins. One centimetre-size, amoeboid titanomagnetite crystal (M5, Fig. 3m), present in the most xenolith-rich mugearite, may also be a xenocryst because it is higher in Al2O3 and TiO2 than oxide crystals in the host lavas or anorthosite xenoliths, suggesting that it crystallized from a more evolved magma composition. One centimetre-size, clear, compositionally zoned clinopyroxene megacryst (M6, Fig. 3n) has a core low in TiO2 and Al2O3 (0·75 wt% and 5·5 wt%, respectively, Fig. 6b and c) comparable with the pyroxene in the amphibole-free gabbros. From core to rim, the crystal becomes progressively higher in Al2O3 over a narrow range in TiO2. Other complexly zoned, lavender-coloured clinopyroxene crystals and plagioclase are abundant in most lava samples and may be disaggregated from gabbro xenoliths like Gn2 and Gn3. Because their compositions overlap with those of the host lava (Fig. 6b and c), we do not consider them separately but have included their compositions in the field for host lavas.
|
DISCUSSION |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Origin of xenoliths
Despite its volumetric significance, direct knowledge of the composition of the lower ocean crust and underlying mantle comes from relatively few studies of rocks exposed in fracture zones and a few Ocean Drilling Program (ODP) drill sites. As our interpretation of the origin of the xenoliths relies primarily on mineral compositions and rock textures, we summarize pertinent data from the Hess Deep (Hekinian et al., 1993
; Allan & Dick, 1996) and Mid-Cayman Rise (Elthon, 1987) because these studies provide detailed mineral chemistry. Fields for these minerals are shown for comparison with our xenolith and xenocryst compositions in Figs 5–8
. The Hess Deep drill site 895 provides a view into the lower crust and upper mantle under a fast-spreading center, whereas the Cayman Rise study provides a view into magma chambers beneath a slow-spreading and dying ridge segment, presumably analogous to the abandoned spreading center under Davidson Seamount. At Hess Deep, the stratigraphic section extends from MORB, through diabase and isotropic gabbro, into gabbroic cumulates and mantle rocks of lherzolite and harzburgite (e.g. Hekinian et al. 1993; Allan & Dick, 1996; Dick & Natland, 1996). Diagnostic mineral compositions of the mantle rocks are highly magnesian olivine (>Fo90) with high NiO and low CaO contents (Fig. 5), clino- and orthopyroxene and spinel with high Mg-numbers. Both spinel and clinopyroxene have low TiO2 and relatively high Cr2O3 contents (Figs 6 and 7). Although some of the Hess Deep ultramafic and mafic rocks have undergone complex wall-rock–melt interaction, olivine with >Fo90, high NiO and low CaO contents in dunite is undoubtedly of mantle origin. Olivine in gabbro cumulates extends to lower Fo compositions but overlaps with mantle olivine in the higher Fo range. However, CaO in these olivines is consistently lower than in phenocrysts in ocean floor basalt, reflecting slow cooling and/or greater pressure. Spinel compositions in cumulate gabbros are distinctly lower in Mg-number and have several times greater TiO2 contents than those of mantle rocks.
No mantle rocks were recovered at the Mid-Cayman Rise (Elthon, 1987) but the suite of gabbros, troctolite, and anorthosites described show the diversity of magma cumulates present under a slow-spreading center. The olivine in these rocks spans a large range in Fo (88–73) and NiO contents (0·26–0·11 wt%) but all have low CaO contents (0·03–0·08 wt%), reflecting slow cooling of deep-seated rocks. Likewise, clinopyroxenes have Mg-numbers ranging from 88 to 62, with TiO2 contents higher than for any of the mantle rocks but typical for tholeiitic compositions. Spinel compositions have an enormous range (Mg-numbers 60–10) presumably as a result of re-equilibration at lower temperatures (Elthon, 1987). Plagioclase compositions ranges from >An70 to An35, but all are low in K2O (0·02–0·22 wt%), typical of normal MORB (N-MORB).
Mantle xenoliths
The mineral compositions of the lherzolite, pyroxenite, and dunite with >Fo90 olivine indicate a mantle origin for these xenoliths. The high Mg-numbers for both clino- and orthopyroxene with low TiO2 and Al2O3 contents (Fig. 6), as well as the low CaO and high NiO contents of olivine (Fig. 5), are comparable with those in harzburgite and spinel lherzolite recovered from beneath the ocean crust in Hess Deep (e.g. Hekinian et al., 1993; Allan & Dick, 1996). Compositions of spinel enclosed in olivine in two of the lherzolites (L1, L2) and in clinopyroxene in one of the pyroxenite (P1) samples are low in TiO2, similar to those from Hess Deep (Fig. 7), although the spinel in the porphyroclastic sample (L1) has a somewhat higher Cr-number. Except for L1, all plot within the field for abyssal peridotites (Dick & Bullen, 1984). The olivine is only slightly strained and has no kink bands except for that in the porphyroclastic lherzolite (L1). The lamellae in the orthopyroxene are undeformed, indicating that no deformation occurred after exsolution. The mantle xenoliths in the seamount lavas appear less deformed than the Hess Deep mantle rocks and, unlike them, show no evidence for retrograde metamorphism in the form of serpentine and greenschist minerals.
The textures and the exceedingly low TiO2 content of clinopyroxene and spinel suggest that these xenoliths are of depleted upper mantle from which N-MORB has been extracted. Glass penetrating into L1 is that of the host magma and not interstitial glass produced by partial melting. Similar mantle xenoliths have been found in alkalic lavas from all over the world, including continental rifts (e.g. Ellis, 1976; McGuire, 1988) and ocean islands such as Hawaii (Sen, 1988), Tahiti (Qi et al., 1994), and the Canary Islands (e.g. Neumann, 1991; Neumann et al., 2000, and references therein). These xenoliths are fragments of upper mantle entrained in ascending alkalic basalt magma.
Ocean crust cumulates
The low-K2O in plagioclase of the troctolite and several gabbro samples (Gn1, Gs1, Gs2, Gs3), as well as the low TiO2 and Al2O3 in clinopyroxene in these gabbros, indicates crystallization from tholeiitic melts (LeBas, 1962; Basaltic Volcanism Study Project, 1981). These gabbros, together with the dunite with olivine of <Fo88, are similar to dunite and gabbro (Figs 5–7) overlying the more depleted mantle peridotite at Hess Deep (e.g. Hekinian et al., 1993; Allan & Dick, 1996), and also closely resemble ocean crust gabbros from the Mid-Cayman (Elthon, 1987) and East Pacific Rise (Hekinian et al., 1985). They are probably ocean crust cumulates from the underlying, abandoned spreading center.
The medium- to coarse-grained, allotriomorphic– to hypidiomorphic–granular textures of the troctolite and gabbro samples are typical of magmatic textures, as are the twinning and zoning of clinopyroxene and the poikilitically enclosed spinel in the olivines of the troctolite and dunites. These spinel compositions are lower in Mg-number and higher in TiO2 than in the lherzolite and pyroxenite (Fig. 7) and suggest an ocean crustal origin. Although the cores of plagioclase with K2O < 0·20 wt% are labradorite with >An55, rims of some crystals at the margins of the xenoliths have lower An and higher K2O overgrowth, probably as a result of reaction with the host lava (Fig. 8). One anorthosite xenolith (A10) has plagioclase that plots along the low-K2O trend of the ocean crust xenoliths (Fig. 8), but the An20–32 compositions suggest that it crystallized from a melt more evolved than typical ocean-ridge basalt (e.g. Basaltic Volcanism Study Project, 1981; Bryan, 1983; Stakes et al., 1984; Davis & Clague, 1987, 1990). Similar, only slightly less sodic plagioclase from inclusions in Mid-Cayman Rise basalts was proposed to have crystallized from the last melt remnants at a dying spreading center (Elthon, 1987). Because Davidson Seamount is built on top of an extinct spreading center (Lonsdale, 1991), this xenolith could be related to the final stage of spreading. Similar gabbro, dunite, and troctolite inclusions found in lavas from ocean islands (e.g. Hawaii, Clague & Chen, 1986; Schmincke et al., 1998; Canary Islands, Neumann et al., 2000) have also been interpreted as ocean crust cumulates.
The compositionally zoned clinopyroxene megacryst (M6, Table 2, Fig. 3n) is probably also an ocean crust fragment. Its low TiO2 and Al2O3 composition indicates crystallization from a tholeiitic parent magma (e.g. LeBas, 1962; Basaltic Volcanism Study Project, 1981; Davis & Clague, 1990). Alternatively, it could be a cumulate from a tholeiitic magma related to seamount formation. Tholeiitic basalt is rare among the seamount lava samples, but one sample was recovered from Pioneer Seamount, and several samples of tholeiitic to transitional basalt were recovered on the deepest dive (T145) at the southernmost end of Davidson Seamount. An Ar–Ar age of 10 Ma (D. A. Clague, unpublished data) for one of these basalts suggests that it is not from an early shield-building stage but is one of the youngest eruptions, postdating the eruption of the xenolith-rich lavas.
Alkalic cumulates related to the host magmatic system
With the exception of anorthosite (A10) mentioned above, the feldspar of the anorthosite xenoliths and of the primary amphibole-gabbros have K2O contents (Fig. 8) consistently higher than those found in ocean crust gabbros, indicating that they crystallized from alkalic magmas. The high K2O and TiO2 contents of the amphibole are also distinct from that of amphibole in MORB gabbros (Fig. 9). Kaersutite and biotite in the gabbros and as megacrysts are diagnostic of a volatile-rich, alkalic melt. Salitic clinopyroxene, high in TiO2 and Al2O3, occurs in three gabbro samples (Gn2, Gn3, Gp5) and is typical of alkalic lava, in contrast to the Cr-diopside of the ocean crust and mantle samples (LeBas, 1962; Basaltic Volcanism Study Project, 1981). Olivine, present only in the two largest amphibole-free gabbro xenoliths (Gn2, Gn3), is reversely zoned with cores lower in Fo and in NiO than the rims or in the host lava. This zoning suggests that they crystallized from a melt more evolved than the host lava. All of these samples exhibit textures that are clearly magmatic, such as large crystals (to 0·9 cm) bounded in part by crystal faces; amphibole poikilitically enclosing euhedral plagioclase, plagioclase enclosing euhedral apatite and/or subhedral magnetite; and twinning and zoning in the clinopyroxene. Although compositional zoning is common in the pyroxene and ubiquitous in the feldspar, the kaersutite is relatively homogeneous for a given sample. All of the kaersutite analyzed, whether in gabbro or occurring as megacrysts, shows a narrower compositional range (Fig. 9) than for amphibole in xenoliths of similar alkalic lavas from the Canary Islands (Neumann et al., 2000) suggesting crystallization from similar alkalic melt compositions. The more sodic, higher K2O cores of plagioclase, and the presence of apatite, titanomagnetite, and especially of zircon inclusions in plagioclase indicate that the xenoliths formed from alkalic melts similar to, but more evolved than, the host lavas. All of these features suggest that these xenoliths are genetically related to the alkalic volcanism that built the main edifices of these seamounts and that they probably crystallized at the margins of the magma reservoirs or along transport paths.
Depth of xenolith origin
Mantle xenoliths
The presence of mantle xenoliths unequivocally indicates that the host magmas originated in the upper mantle. Dense mantle xenoliths, albeit small ones, preclude prolonged storage in shallow reservoirs, where they would have settled out (Clague, 1987). Spinel lherzolite xenoliths demonstrate that the host magma rose from a depth of at least 25 km, below the plagioclase stability field (0·8–1 GPa, e.g. Gasparik, 1984; Sen, 1985). Despite the many thermometers published for spinel lherzolite (e.g. Wood & Banno, 1973; Wells, 1977; Mercier, 1980; Lindsley 1983; Lindsley & Andersen, 1983; Sen, 1985; Brey & Köhler, 1990), there are no reliable geobarometers available. Within and between the various thermometers, there are large uncertainties, but collectively calculated temperatures range from 800 to 1100°C, with most from 900° to 1000°C. Because the two-pyroxene thermometer of Brey & Köhler (1990) is based on the largest dataset, and they provided a comparative assessment of the various other two-pyroxene geothermometers, we used it to calculate temperatures ranging from 1118° to 819°C for the samples with coexisting clino- and orthopyroxene (Table 10). These temperatures are within the range of published values as well as those determined experimentally for the spinel peridotite stability field (900–1100°C, 0·8–1·6 GPa, e.g. Gasparik, 1984; Sen, 1988). Because exsolution lamellae are present, the range we calculated probably represents subsolidus temperatures. Similar temperatures were calculated for spinel lherzolite xenoliths from Hawaii (Sen et al., 2005), the Society Islands (Qi et al., 1994) and the Canary Islands (Neumann, 1991); for the latter two estimated pressures were 1·2–1·6 GPa.
|
The clinopyroxene structural geobarometer of Nimis (1999
) can be applied to a range of pressures and compositions and does not require knowledge of the equilibrium liquid compositions (Nimis & Ulmer, 1998; Nimis, 1999). Using the major element composition of clinopyroxene alone, pressures can be calculated for anhydrous (BA) or hydrous basalt (BH) and for more evolved liquid compositions along the tholeiitic (BT) and mildly alkaline (MA) trends. Except for the BA model, pressures are temperature dependent and small changes in temperature can result in large uncertainties in pressure (0·2 GPa). Using the temperatures determined with the Brey & Köhler thermometer, we obtained pressures ranging from 0·9 to 1·7 GPa (Table 10) with the BT model for the mantle xenoliths. As higher temperatures yield lower pressures, these are maxima.
Ocean crust xenoliths
The xenoliths inferred to be ocean crust cumulates are similar to gabbros from the Mid-Cayman Rise (Elthon, 1987) and to gabbros and some dunites overlying the harzburgite drilled at ODP site 895 near the Hess Deep (Allan & Dick, 1996, Figs 5–8). Elthon (1987) proposed that the Mid-Cayman Rise gabbros did not crystallize from typical low-pressure (1 atm to 0·2 GPa) MORB magmas but probably formed at moderate pressure (0·5–1·0 GPa) within deep-seated magma chambers under this slow-spreading center, after spreading ceased. The ocean crust xenoliths from Davidson, like those from the Cayman Trough, formed within ocean crust layer 3. Pressures, however, must be less than 1· 0 GPa because the xenoliths contain abundant plagioclase. Pressures estimated using geobarometers based on mineral compositions are relatively unconstrained. Pressures calculated arbitrarily at 1000° and 1100°C for clinopyroxene in gabbro and the clinopyroxene megacryst range widely from 0·68 to 1·3 GPa (Table 10), and are greater than our physical estimates presented below. Pyroxene compositions in samples Gs1–Gs3 may not be suitable for these calculations because of secondary amphibole alteration.
These xenoliths presumably crystallized at normal layer 3 crustal depths but were re-equilibrated to slightly greater pressure as the seamount grew and depressed the crust. We can therefore estimate the depth of equilibration of the ocean crust xenoliths by determining the depth to the crust–mantle boundary beneath the seamounts. Miller et al. (1992) have shown that complex tectonics resulted in local thickening of underplated oceanic crust at the central California margin but farther offshore, near Davidson Seamount, they show crustal thickness of 7 km. Adding about 2 km for the height of the seamounts to the crustal thickness and an additional 2–5 km estimated for isostatic depression as a result of the seamount load, we infer that the crust–mantle boundary was about 11–14 km deep (equivalent to <0·5 GPa) when the seamounts formed.
Alkalic cumulates
All of the mineral compositions of the xenoliths and megacrysts of alkalic affinity indicate that they crystallized from a melt similar to but more evolved than their hawaiite and mugearite host lavas. However, the parental alkalic melts may have crystallized at depths greater or less than the depths of origin of the mantle and crustal xenoliths. We have used a variety of techniques to estimate the temperature and pressure at which the alkalic cumulate xenoliths could have crystallized from the seamount magmas.
Kaersutitic amphibole can crystallize over a considerable pressure range (e.g. Best, 1970; Dawson & Smith, 1982, and references therein) and is not diagnostic of a specific depth. It is unstable and tends to oxidize at magmatic temperatures in shallow reservoirs and is absent at pressures above the amphibole stability field (>2·5 GPa, e.g. Niida & Green, 1999). Similar kaersutite inclusions and megacrysts from continental settings in Arizona (Best, 1975), Australia (Ellis, 1976), and the Tertiary volcanic rocks of Germany (Vinx & Jung, 1977) were interpreted as having formed in the upper mantle and/or lower crust at >1 GPa. The amphibole compositions in the gabbros and of megacrysts indicate magmatic temperatures of >1100°C at 0·3 GPa, using the semi-empirical thermometer of Otten (1984). At 0·5 GPa the temperatures are about 80–90° lower and are in agreement with the ilmenite–magnetite temperature calculated for sample Gp8 (Table 10). Only one sample (Gp5) contains clinopyroxene, which apparently crystallized from a mildly alkalic melt, based on the Ti and Al contents. Using the temperature obtained for coexisting magnetite–ilmenite (Andersen et al., 1983), pressures calculated with the Nimis clinopyroxene barometer range from 0·49 to 0·55 GPa for the anhydrous (BA) and from 0·8 to 0·9 GPa for the alkaline (MA) model (Table 10). The MA model always yields pressures significantly higher than the BA model. Considering that the minerals in sample Gp8 span a narrow compositional range and that the magnetite–ilmenite thermometer is probably one of the most reliable ones, this calculation may provide the best constraint on pressure. Similar pressures (Table 10) were calculated with the Al-in-hornblende barometer of Schmidt (1992), although it may not be applicable to these compositions as it was calibrated for amphibole crystallized from tonalite.
The MELTS program (Ghiorso & Sack, 1995) has been widely used to evaluate liquid lines of descent for various magmas. Davidson lava compositions are not well replicated at any single pressure from 0·1 to 0·9 GPa (Fig. 10) with a range of water contents (0·5–1%) and oxygen fugacity at the quartz–fayalite–magnetite buffer. We used the high- and low-SiO2 end-members in our modeling. The observed range of lava and glass compositions clearly requires a range of starting parental magma compositions. The prominent decrease in CaO/Al2O3 observed in the lava and glass compositions (Fig. 10a) requires significant clinopyroxene fractionation. Clinopyroxene becomes the dominant phase with increasing pressure, resulting in a large increase in alkalis without changing the silica content significantly in the basalt to hawaiite range, as observed for the Davidson lavas. However, at pressures of 0·7–0·9 GPa at low MgO contents, amphibole never appears as a crystallizing phase and the SiO2 enrichment observed in trachyte cannot be attained (Fig. 10b). The required SiO2 enrichment is possible at low pressure (0·1–0·3 GPa), but amphibole is again absent from the crystallizing assemblage and the Al2O3 enrichment observed in the trachyte is not attained. We could not find a combination of pressure and water content that produced both the SiO2 and Al2O3 enrichment observed in trachyte, and none of the runs produced amphibole, clearly an important phase in the evolution of these lavas. Crystallization of kaersutite at the expense of plagioclase and garnet (not observed) would result in higher SiO2 and Al2O3 in the melt and match the observed trachyte compositions. In summary, the MELTS program does not match the observed mineralogy of the xenoliths or the sequence of observed resultant liquid compositions at any pressure or combination of pressures. Our evaluation of its utility in modeling hydrous alkalic basalt compositions suggests that the recognized problem in modeling intermediate to silicic calc-alkaline compositions involving fractionation of hornblende and biotite (Ghiorso & Sack, 1995) extends to more alkaline compositions as well.
|
Experimental studies of alkalic lavas similar to the host lavas may provide better constraints for their evolution. Nevkasil et al. (2004
) showed that an increase in alkalis relative to nearly constant silica, as seen in Davidson lavas and commonly observed in alkalic lava suites, is due to the dominance of clinopyroxene and suppression of plagioclase in early fractionating assemblages at elevated pressures (0·9 GPa, 1100°C). At the same pressure, but at intermediate temperatures (1090–940°C), they found that kaersutite became a dominant phase under hydrous (0·5% bulk water) conditions but was replaced by a Ti-rich biotite under less hydrous conditions and at the lower end of the temperature range. The higher temperature (>1000°C) calculated for the amphibole-gabbro (Table 10) is close to that in the high-pressure experiments of Nevkasil et al. (2004) that yielded kaersutite, apatite, and ilmenite. The compositions of these experimentally derived phases, including plagioclase, are also similar to those in Davidson xenoliths (Fig. 9). In agreement with earlier experimental studies (Mahood & Baker, 1986), Nevkasil et al. (2004) confirmed that plagioclase tends to be more sodic at elevated pressures. Experimental studies of MORB at high pressures (Bender et al., 1978; Green et al., 1979) also showed a decrease of An contents in plagioclase with increasing pressure, suggesting that this effect may be independent of lava composition. The similarity of the Davidson xenoliths to these experimental results suggests that the Davidson magmas crystallized the alkalic cumulates at pressures <0·9 GPa, or a depth of about 24 km, at the base of the lithosphere (Zhang & Lay, 1999).
Nevkasil et al. (2004) further suggested that water content could play a more important role than pressure, especially as a variable in suppressing early feldspar crystallization. Bulk water contents measured for some hawaiite glasses from Davidson are 0·7 wt% (Davis & Clague, 2003), but must have been much greater at the margins of the magma reservoirs where the amphibole crystallized. The high fluorine and chlorine contents of the amphibole and apatite (1·4 and 0·9 wt%, respectively) indicate high halogen contents in the magmas, but their effects on crystallization have not been well investigated experimentally and so cannot be evaluated.
Transport of xenoliths to the surface
We have presented evidence that these alkalic magmas fractionated at 0·7–0·9 GPa, at or near the base of the lithosphere, before migrating to the surface and eruption. During their ascent they entrained and transported some of the partly crystallized wall rocks (alkalic cumulates and megacrysts) of their deep magma reservoirs and crystalline xenoliths of upper mantle (<0·9 GPa) and ocean crustal rocks (<0·5 GPa). These magmas probably rose to the surface rapidly to maintain the dense xenoliths in suspension, and because of the high volatile contents we infer for the parent magmas and the explosive character of many of the eruptions (Davis & Clague, 2003).
The kaersutite, biotite, and apatite, all with high fluorine and chlorine contents, testify to high contents of water and halogens in the magma. We think these volatiles accumulated and became enriched at the upper margins of the alkalic melt pockets and in veins extending into the mantle country rock and crystallized hydrous phases such as kaersutite and biotite. With continued fractionation, the bulk water contents increased, thereby lowering the melt density and viscosity, and the increase in volatiles pressurized the system and eventually propelled the lava to the eruption site on the sea floor. Many of the xenoliths and megacrysts are contained in highly vesicular hyaloclastite breccias, demonstrating the explosivity of the eruptions (Davis & Clague, 2003).
The rapid ascent rate of the host magma is supported by the lack of diffusion of CaO in unaltered olivine (Klügel, 1998), by thin or absent lava selvages on mantle and some ocean crust xenoliths, and by decompression that fractured the xenoliths and also increased the temperature of the host lava. The increased temperatures caused extensive resorption of sodic plagioclase and the formation of thin overgrowths of more calcic plagioclase. The rounded shapes of amphibole and embayed margins of titanomagnetite megacrysts suggest dissolution and ablation of these lower temperature phases. Many of the fractures along congugate joint surfaces that were filled by host lava probably originated at this time as a result of rapid decompression.
Individual eruptions might have been initiated by, or at least were aided by, regional extensional tectonics. Synchronous volcanism occurred at geographically widely separated places on- and offshore along the continental margin during the Miocene that might have been related to movement along major faults as the tectonic regime changed from a convergent to a transform margin (e.g. Davis et al., 1995, 2002; Dickinson, 1997). Bailey (1970, 1972) suggested that lithospheric structures focused areas of alkalic volcanism in continental rift settings. The alignment of volcanic cones at Davidson Seamount parallel to the ocean crust fabric suggests a similar pattern with ascending melts channeled along existing zones of weakness.
|
SUMMARY AND CONCLUSIONS |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
The three types of xenoliths included in the seamount lavas provide information on the characteristics of the mantle and crust underlying the volcanoes. The mantle xenoliths indicate that the initial melt rose from mantle depths below the plagioclase stability field (
1 GPa). The ocean crust xenoliths are cumulates from the final stages of spreading centers that were abandoned when the tectonic regime changed from a divergent to a transform margin. The alkalic gabbros, anorthosites, and kaersutite and titanomagnetite megacrysts are cumulates formed at moderate pressure (0·5–0·9 GPa) at the base of the lithosphere from melts that are genetically related to, but more evolved than, the host lava. Increasing water and other volatile constituents decreased magma density and pressurized the system, leading to rapid ascent and eruption on the ocean floor. |
SUPPLEMENTARY DATA |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Supplementary data for this paper are available at Journal of Petrology online.
|
ACKNOWLEDGEMENTS |
|---|
Most of the xenolith samples were collected in dredges carried out by Gary Greene and Brent Dalrymple on several USGS cruises in the late 1970s. Four xenoliths are from dives T141 and T142 by Peter Lonsdale and Pat Castillo and three additional xenoliths were collected by Andrew DeVogeleare on dives T426 and T427, funded by NOAA's Ocean Exploration Program. These principal investigators kindly made the additional samples available to us, thereby increasing the number of samples available. We thank the ROV Tiburon pilots and the captain and crew of the Western Flyer for their skill in recovering the dive samples. Robert Oscarson assisted with microprobe analysis. We thank Gautam Sen and Paolo Nimis for sharing their EXCEL spreadsheets for geothermobarometry. Reviews by M. Coombs, R. Fodor, S. Keshav, and especially Editor W. Bohrson of an earlier version greatly improved the manuscript. The support of the David and Lucile Packard Foundation through a grant to MBARI is gratefully acknowledged.
*Corresponding author. Telephone: 831-775-1857. davisa{at}mbari.org
|
REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSHOST LAVADESCRIPTION OF XENOLITHS AND...DISCUSSIONSUMMARY AND CONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
|---|
Allan JF and Dick HJ. (1996) Cr-rich spinel as a tracer for melt migration and melt–wall rock interaction in the mantle: Hess Deep, leg 147. In Mevel C, Gillis KM, Allan JF, Meyer PS (Eds.). Proceedings of the Ocean Drilling Program, Scientific Results, 147(Ocean Drilling Program, College Station, TX) pp. 157–172.
Andersen DJ, Lindsley DH, Davidson PM. (1993) QUILF: a Pascal program to assess equilibria among Fe–Mg–Mn–Ti oxides, pyroxenes, olivine, and quartz. Computers and Geosciences 19:91333–1350.
Bailey DK. (1970) Volatile flux, heat focusing and the generation of magma. Geologisches Jahrbuch, Special Issue 2:177–186.
Bailey DK. (1972) Uplift, rifting and magmatism in continental plates. Journal of Earth Sciences 8:225–239.
. Basaltic Volcanism Study Project. (1981) Basaltic Volcanism on the Terrestrial Planets(Pergamon Press, New York) pp. 1288.
Bender JF, Hodges FN, Bence AE. (1978) Petrogenesis of basalts from the project FAMOUS area: experimental studies from 0 to 15 kilobars. Earth and Planetary Science Letters 41:277–302.[CrossRef][Web of Science]
Best MG. (1975) Mantle-derived amphibole within inclusions and kindred megacrysts in basanitic lavas, Grand Canyon, Arizona. Contributions to Mineralogy and Petrology 27:25–45.
Bohrson WA and Clague DA. (1988) Origin of ultramafic xenoliths containing exsolved pyroxenes from Hualalai volcano, Hawaii. Contributions to Mineralogy and Petrology 100:139–155.[CrossRef][Web of Science]
Brey GP and Köhler T. (1990) Geothermobarometry in four-phase lherzolites II. New thermobarometers and practical assessment of existing thermobarometers. Journal of Petrology 31:1353–1378.
Bryan WB. (1983) Systematics of modal phenocryst assemblages in submarine basalts: petrologic implications. Contributions to Mineralogy and Petrology 83:62–74.[CrossRef][Web of Science]
Cannat M, Chatin F, Whitechurch H, Ceuleneer G. (1997) Gabbroic rocks trapped in the upper mantle at the Mid-Atlantic Ridge. In Karson JA, Cannat M, Miller DJ, Elthon D (Eds.). Proceedings of the Ocean Drilling Program, Scientific Results, 153(Ocean Drilling Program, College Station, TX) pp. 243–264.
Clague DA. (1987) Hawaiian xenolith populations, magma supply rates and development of magma chambers. Bulletin of Volcanology 49:4577–587.[CrossRef]
Clague DA and Chen C.-H. (1986) Ocean crust xenoliths from Hualalai volcano, Hawaii. Geological Society of America, Abstracts 18:565.
Cox KG, Bell JD, Pankhurst RJ. (1979) The Interpretation of Igneous Rocks(Allen and Unwin, London) pp. 450.
Davis AS and Clague DA. (1987) Geochemistry, mineralogy, and petrogenesis of basalt from the Gorda Ridge. Journal of Geophysical Research 92:10467–10483.
Davis AS and Clague DA. (1990) Gabbroic xenoliths from the northern Gorda Ridge: implications for magma chamber processes under slow-spreading centers. Journal of Geophysical Research 95:10885–10990.
Davis AS and Clague DA. (2003) Hyaloclastite from Miocene seamounts offshore central California: compositions, eruption styles, and depositional processes. In White JDL, Smellie JL, Clague DA (Eds.). Explosive Subaqueous Volcanism. Geophysical Monograph, American Geophysical Union 140: pp. 129–142.
Davis AS, Clague DA, Friesen WB. (1994) Petrology and mineral chemistry of basalt from Escanaba Trough, southern Gorda Ridge. US Geological Survey Bulletin 2022:153–170.
Davis AS, Gunn SH, Bohrson WA, Gray LB, Hein JR. (1995) Chemically diverse sporadic volcanism at seamounts offshore southern and Baja California. Geological Society of America Bulletin 107:554–570.
Davis AS, Clague DA, Bohrson WA, Dalrymple GB, Greene HG. (2002) Seamounts at the continental margin of California: a different kind of oceanic volcanism. Geological Society of America Bulletin 114:316–333.
Dawson JB and Smith JV. (1982) Upper-mantle amphiboles: a review. Mineralogical Magazine 45:35–36.[Web of Science]
Dick H. J. B and Bullen T. (1984) Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatially associated lavas. Contributions to Mineralogy and Petrology 86:54–76.[CrossRef][Web of Science]
Dick HJB and Natland JH. (1996) Late-stage melt evolution and transport in the shallow mantle beneath the East Pacific Rise. In Mevel C, Gillis KM, Allan JF, Meyer PS (Eds.). Proceedings of the Ocean Drilling Program, Scientific Results, 147(Ocean Drilling Program, College Station, TX) pp. 103–134.
Dickinson WR. (1997) Tectonic implications of Cenozoic volcanism in coastal California. Geological Society of America Bulletin 109:936–954.
Dixon JE, Clague DA, Eissen JP. (1986) Gabbroic xenoliths and host ferrobasalt from the southern Juan de Fuca Ridge. Journal of Geophysical Research 91:3795–3820.
Ellis DJ. (1976) High pressure cognate inclusions in the Newer Volcanics of Victoria. Contributions to Mineralogy and Petrology 58:149–180.[CrossRef][Web of Science]
Elthon D. (1987) Petrology of gabbroic rocks from the mid-Cayman rise spreading center. Journal of Geophysical Research 92:658–682.
Fodor RV and Galar P. (1997) A view into the subsurface of Mauna Kea volcano, Hawaii: crystallization processes interpreted through the petrology and petrography of gabbroic and ultramafic xenoliths. Journal of Petrology 38:5581–624.[CrossRef][Web of Science]
Fodor RV and Moore RB. (1994) Petrology of gabbroic xenoliths in 1960 Kilauea basalt: crystalline remnants of prior (1955) magmatism. Bulletin of Volcanology 56:62–74.[CrossRef][Web of Science]
Fodor RV and Vandermeyden HJ. (1988) Petrology of gabbroic xenoliths from Mauna Kea volcano, Hawaii. Journal of Geophysical Research 93:B54435–4452.
Frey FA and Prinz M. (1978) Ultramafic inclusions from San Carlos, Arizona: petrologic and geochemical data bearing on their petrogenesis. Earth and Planetary Science Letters 38:129–176.[CrossRef][Web of Science]
Gaggero L and Cortesogno L. (1997) Data report: metamorphic mineralogy of Leg 153 gabbro. In Karson JA, Cannat M, Miller DJ, Elthon D (Eds.). Proceedings of the Ocean Drilling Program, Scientific Results, 153(Ocean Drilling Program, College Station, TX) pp. 531–546.
Gasparik T. (1984) Two-pyroxene thermobarometry with new experimental data in the system CaO–MgO–Al2O3–SiO2. Contributions to Mineralogy and Petrology 87:87–97.[CrossRef][Web of Science]
Ghiorso MS and Sack RO. (1995) Chemical mass transfer in magmatic processes IV: A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology 119:197–212.[Web of Science]
Green DH, Hibberson WO, Jaques AL. (1979) Petrogenesis of mid-ocean ridge basalts. In McElhinny MW (Ed.). The Earth: its Origin, Structure and Evolution(Academic Press, London) pp. 265–299.
Hekinian R, Hebert R, Maury RC, Berger ET. (1985) Orthopyroxene-bearing xenoliths in basalts from the East Pacific Rise axis near 12°50'N. Bulletin de Minéralogie 108:691–698.
Hekinian R, Bideau D, Francheteau , J., Cheminee JL, Armijo R, Lonsdale P, Blum N. (1993) Petrology of the East Pacific Rise and upper mantle exposed in Hess Deep (Eastern equatorial Pacific). Journal of Geophysical Research 98:8069–8094.
Kempton PD. (1987) Mineralogical and geochemical evidence for differing styles of metasomatism in spinel lherzolite xenoliths: enriched mantle source regions of basalts? In Menzies MA and Hawkesworth CJ (Eds.). Mantle Metasomatism(Academic Press, London) pp. 45–87.
Klügel A. (1998) Reactions between mantle xenoliths and host magma beneath La Palma (Canary Islands): constraints on magma ascent rates and crustal reservoirs. Contributions to Mineralogy and Petrology 131:237–257.[CrossRef][Web of Science]
Leake BE. (1978) Nomenclature of amphiboles. Canadian Mineralogist 16:501–520.
LeBas MJ. (1962) The role of aluminum in igneous clinopyroxene with relation to their parentage. American Journal of Science 260:267–288.
Lindsley DH. (1983) Pyroxene thermometry. American Mineralogist 68:477–493.[Abstract]
Lindsley DH and Andersen DJ. (1983) A two-pyroxene thermometer. Proceedings of the 13th Lunar and Planetary Conference. Journal of Geophysical Research, Supplement 88: pp. A887–A906.
Lonsdale P. (1991) Structural patterns of the Pacific floor offshore of peninsular California. In Dauphin JP, Ness G, Simoneit BRT (Eds.). The Gulf and Peninsular Province of the Californias. American Association of Petroleum Geologists, Memoirs 47: pp. 87–125.
Mahood GA and Baker DR. (1986) Experimental constraints on depths of fractionation of mildly alkalic basalts and associated felsic rocks: Pantelleria, Strait of Sicily. Contributions to Mineralogy and Petrology 93:2251–264.[CrossRef][Web of Science]
McGuire AV. (1988) Petrology of mantle xenoliths from Harrat al Kishb: the mantle beneath western Saudi Arabia. Journal of Petrology 29:73–92.
Mercier J.-C. C. (1980) Single-pyroxene thermobarometry. Tectonophysics 70:1–37.[CrossRef][Web of Science]
Miller KC, Howie JM, Ruppert SD. (1992) Shortening within underplated oceanic crust beneath the central California margin. Journal of Geophysical Research 97:19961–19980.
Nekvasil H, Dondolini A, Horn J, Filiberto J, Long H, Lindsley DH. (2004) The origin and evolution of silica-saturated alkalic suites: an experimental study. Journal of Petrology 45:4693–721.
Neumann E.-R. (1991) Ultramafic and mafic xenoliths from Hierro, Canary Islands: evidence for melt infiltration in the upper mantle. Contributions to Mineralogy and Petrology 106:236–252.[CrossRef][Web of Science]
Neuman E.-R, Sorensen VB, Simonsen SL, Johnson K. (2000) Gabbroic xenoliths from La Palma, Tenerife and Lanzarote, Canary Islands: evidence for reactions between mafic alkaline Canary Islands melts and old oceanic crust. Journal of Volcanology and Geothermal Research 103:313–342.[CrossRef][Web of Science]
Niida K and Green DH. (1999) Stability and chemical composition of pargasitic amphibole in MORB pyrolite under upper mantle conditions. Contributions to Mineralogy and Petrology 135:18–40.[CrossRef][Web of Science]
Nimis P. (1999) Clinopyroxene geobarometry of magmatic rocks, Part 2: Structural geobarometers for basic to acid, tholeiitic and mildly alkaline magmatic systems. Contributions to Mineralogy and Petrology 135:162–74.[CrossRef][Web of Science]
Nimis P and Ulmer P. (1998) Clinopyroxene geobarometry of magmatic rocks, Part 1: An expanded structural geobarometer for anhydrous and hydrous, basic and ultrabasic systems. Contributions to Mineralogy and Petrology 133:1–2122–135.[CrossRef][Web of Science]
Otten MT. (1984) The origin of brown hornblende in the Artfjället gabbro and dolerites. Contributions to Mineralogy and Petrology 86:189–199.[CrossRef][Web of Science]
Paduan JB, Clague DA, Davis AS. (2004) Evidence that three seamounts off Southern California were ancient islands. EOS Transactions, American Geophysical Union 85:471463.
Pike J. E. N and Schwarzman EC. (1977) Classification of textures in ultramafic xenoliths. Journal of Geology 85:49–61.[Web of Science]
Qi Q, Beard BL, Jin Y, Taylor LA. (1994) Petrology and geochemistry of Al-augite and Cr-diopside group mantle xenoliths from Tahiti, Society Islands. International Geology Review 36:152–178.
Schmidt MW. (1992) Amphibole composition in tonalite as a function of pressure: an experimental calibration of the Al-in-hornblende barometer. Contributions to Mineralogy and Petrology 110:304–310.[CrossRef][Web of Science]
Schmincke HU, Kluegel A, Hansteen TH, Hoernle K, van den Bogaard P. (1998) Samples from the ocean crust beneath Gran Canaria, La Palma and Lanzarote (Canary Islands). Earth and Planetary Science Letters 163:342–360.
Sen G. (1985) Experimental determination of pyroxene compositions in the system CaO–MgO–Al2O3–SiO2 at 900–1200°C and 10–15 kbar using PbO and H2O fluxes. American Mineralogist 70:678–695.[Abstract]
Sen G. (1988) Petrogenesis of spinel lherzolite and pyroxenite suite xenoliths from the Koolau shield, Oahu, Hawaii: implications for petrology of the post-eruptive lithosphere beneath Oahu. Contributions to Mineralogy and Petrology 100:61–91.[CrossRef][Web of Science]
Sen G and Presnall DC. (1986) Petrogenesis of dunite xenoliths from Koolau Volcano, Hawaii: implications for Hawaiian volcanism. Journal of Petrology 27:197–217.
Sen G, Keshav S, Bizimis M. (2005) Hawaiian xenoliths and magmas: composition and thermal character of the lithosphere. American Mineralogist 90:871–887.
Stakes DS, Shervais JW, Hopson CA. (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 89:NB86995–7028.
Takahashi E. (1980) Thermal history of lherzolite xenoliths—I. Petrology of lherzolite xenoliths from the Ichinomegata crater, Oga peninsula, northeast Japan. Geochimica et Cosmochimica Acta 44:1643–1658.[CrossRef][Web of Science]
Vinx R and Jung D. (1977) Pargasitic–kaersutitic amphibole from a basanitic diatreme at the Rosenberg, North Kassel (North Germany). Contributions to Mineralogy and Petrology 65:135–142.[CrossRef][Web of Science]
Wells RA. (1977) Pyroxene thermometry in simple and complex systems. Contributions to Mineralogy and Petrology 62:129–139.[CrossRef][Web of Science]
Wood BJ and Banno S. (1973) Garnet–orthopyroxene–clinopyroxene relationships in simple and complex systems. Contributions to Mineralogy and Petrology 42:109–124.[CrossRef][Web of Science]
Zhang Y.-S and Lay T. (1999) Evolution of oceanic upper mantle structure. Physics of the Earth and Planetary Interiors 114:71–80.[CrossRef][Web of Science]
CiteULike 
Connotea 
Del.icio.us What's this?
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||