suprasubduction-zone ophiolites: is there really an ... - my articles/gsa-sp438-2008... ·...

191

Metcalf, R.V., and Shervais, J.W., 2008, Suprasubduction-zone ophiolites: Is there really an ophiolite conundrum?, in Wright, J.E., and Shervais, J.W., eds., Ophiolites, Arcs, and Batholiths: A Tribute to Cliff Hopson: Geological Society of America Special Paper 438, p. 191–222, doi: 10.1130/2008.2438(07). For per-mission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

The Geological Society of AmericaSpecial Paper 438

2008

Suprasubduction-zone ophiolites: Is there really an ophiolite conundrum?

Rodney V. Metcalf†

Department of Geoscience, University of Nevada–Las Vegas, Las Vegas, Nevada 89154-4010, USA

John W. Shervais‡

Department of Geology, Utah State University, Logan, Utah 84322-4505, USA

ABSTRACT

Suprasubduction-zone ophiolites have been recognized in the geologic record for over thirty years. These ophiolites are essentially intact structurally and stratigraphi-cally, show evidence for synmagmatic extension, and contain lavas with geochemical characteristics of arc-volcanic rocks. They are now inferred to have formed by hinge retreat in the forearc of nascent or reconfi gured island arcs. Emplacement of these forearc assemblages onto the leading edge of partially subducted continental margins is a normal part of their evolution. A recent paper has challenged this interpretation. The authors assert that the “ophiolite conundrum” (seafl oor spreading shown by dike complexes versus arc geochemistry) can be resolved by a model called “historical contingency,” which holds that most ophiolites form at mid-ocean ridges that tap upper-mantle sources previously modifi ed by subduction. They support this model with examples of modern mid-ocean ridges where suprasubduction zone–like compo-sitions have been detected (e.g., ridge-trench triple junctions).

The historical contingency model is fl awed for several reasons: (1) the major- and trace-element compositions of magmatic rocks in suprasubduction-zone ophio-lites strongly resemble rocks formed in primitive island-arc settings and exhibit distinct differences from rocks formed at mid-ocean-ridge spreading centers; (2) slab-infl uenced compositions reported from modern ridge-trench triple junctions and subduction reversals are subtle and/or do not compare favorably with either modern subduction zones or suprasubduction-zone ophiolites; (3) crystallization sequences, hydrous minerals, miarolitic cavities, and reaction textures in suprasubduction-zone ophiolites imply crystallization from magmas with high water activities, rather than mid-ocean-ridge systems; (4) models of whole Earth convection, subduction recycling, and ocean-island basalt isotopic compositions ignore the fact that these components represent the residue of slab melting, not the low fi eld strength element–enriched component found in active arc-volcanic suites and suprasubduction-zone ophiolites; and (5) isotopic components indicative of mantle heterogeneities (related to subduc-

†E-mail: [email protected].‡E-mail: [email protected].

192 Metcalf and Shervais

INTRODUCTION

Ophiolites are distinct assemblages of ultramafi c, mafi c, and felsic igneous rocks, commonly associated with siliceous pelagic sediments (cherts), that have long been recognized as important components of mountain belts worldwide (Steinmann, 1906; Hess, 1955). In the 1960s, this assemblage was proposed to represent oceanic crust formed at mid-oceanic spreading centers, a concept that became central to the new theory of plate tectonics (Gass, 1968). A compelling aspect of this proposal was the recognition of sheeted dike complexes in some ophiolites that implied formation by 100% extension (e.g., Troodos; Oman), consistent with the new concept of seafl oor spreading in ocean basins (Moores and Vine, 1971). Dedicated campaigns of deep-ocean drilling, dredging, and seismic-refraction surveys confi rmed the similarity of oceanic crust to ophiolites, although there were differences in detail. As a result, this paradigm became entrenched within the scientifi c com-munity—especially among those who did not work on ophiolites.

Suprasubduction-zone ophiolites have been recognized in the geologic record for over three decades (Miyashiro, 1973; Pearce et al., 1984; Shervais and Kimbrough, 1985). These ophiolites are made up of plutonic rocks and lavas with the mineralogical and geochemical characteristics of arc-plutonic and arc-volcanic rocks, and they are petrologically and chemically distinct from igneous rocks formed at modern spreading centers in the major ocean basins. In general, suprasubduction-zone ophiolites are intact structurally and stratigraphically and show evidence for nearly 100% extension. Such ophiolites are now inferred to have formed primarily by hinge retreat in the forearc of nascent or reconfi gured island arcs, a model derived from studies of Cenozoic subduc-tion systems in the western Pacifi c (Fig. 1; Hawkins et al., 1984; Stern and Bloomer, 1992; Bloomer et al., 1995; Hawkins, 2003). Emplacement of these forearc assemblages onto the leading edge of partially subducted continental margins (Tethyan ophiolites) or exposure by accretionary uplift along an active plate margin (Cordilleran ophiolites) is a normal part of their evolution (e.g., Shervais, 2001). Several recent papers have discussed the develop-ment of suprasubduction-zone ophiolite models, their genesis, and tectonic implications, most notably Shervais (2001), Dilek (2003), Pearce (2003), Hawkins (2003), and Flower (2003).

Moores et al. (2000) challenged the suprasubduction interpre-tation of ophiolite genesis. These authors assert that the “ophio-lite conundrum” (seafl oor spreading shown by dike complexes versus arc geochemistry) can be resolved by a model called “his-torical contingency,” which holds that most ophiolites are formed at mid-ocean ridges that tap upper-mantle sources previously modifi ed by subduction. They support this model with examples of subduction-zone reversal, which place oceanic spreading centers above lithosphere previously modifi ed by subduction (i.e., the Woodlark basin), with examples of modern mid-ocean ridges where suprasubduction zone–like compositions have been detected (e.g., ridge-trench-trench triple junctions), with models of mantle convection that show recycling of oceanic lithosphere on grand scale, and with a discussion of the isotopic components found in ocean-island basalts (OIBs) (Moores et al., 2000).

Moores et al. (2000) also suggest that differences observed in the structural preservation of ophiolites result from distinct spreading environments, not from their subsequent emplace-ment. Thus, ocean crust and ophiolites formed at slow spread-ing centers are highly faulted and commonly have volcanic rocks juxtaposed against serpentine, whereas ocean crust and ophio-lites formed at fast spreading centers tend to be stratigraphically intact and lack the extreme structural attenuation found in slow spreading ocean crust (Moores et al., 2000). Examples of slow spreading ophiolites would include those in the Western Mediterranean (Apennines); examples of fast spreading ophio-lites would include Troodos and Oman.

We suggest that the historical contingency model is fl awed for several reasons: (1) the major- and trace-element composi-tions of magmatic rocks in suprasubduction-zone ophiolites strongly resemble rocks formed in primitive island-arc settings and exhibit distinct, consistent differences from rocks formed at mid-ocean-ridge spreading centers; (2) slab-infl uenced composi-tions reported from modern ridge-trench triple junctions and sub-duction reversals are subtle and/or do not compare favorably with either modern subduction zones or suprasubduction-zone ophio-lites; (3) crystallization sequences, hydrous minerals (hornblende), miarolitic cavities, and reaction textures in suprasubduction-zone ophiolites imply crystallization from magmas with high water activities, rather than mid-ocean-ridge magmatic systems; (4)

tion recycling) are observed in modern mid-ocean-ridge basalts (MORB), but, in con-trast to the prediction of the historical contingency model, these basalts do not exhibit suprasubduction zone–like geochemistry. The formation of suprasubduction-zone ophiolites in the upper plate of subduction zones favors intact preservation either by obduction onto a passive continental margin, or by accretionary uplift above a sub-duction zone. Ophiolites characterized by lavas with MORB geochemistry are typi-cally disrupted and found as fragments in accretionary complexes (e.g., Franciscan), in contrast to suprasubduction-zone ophiolites. This must result from the fact that oceanic crust is unlikely to be obducted for mechanical reasons, but it may be pre-served where it is scraped off of the subducting slab.

Keywords: ophiolite, suprasubduction zone, mid-ocean ridge, geochemistry mantle.

Suprasubduction-zone ophiolites: Is there really an ophiolite conundrum? 193

many suprasubduction-zone ophiolites are overlain by evolved lavas and volcaniclastic rocks that typically are not counted as part of the ophiolite, (5) models of whole Earth convection, sub-duction recycling, and OIB isotopic compositions ignore the fact that these components represent the residue of slab melting, not

the low fi eld strength element–enriched component found in active arc-volcanic suites and suprasubduction-zone ophiolites; and (6) the isotopic components are indicative of mantle hetero-geneities (related to subduction recycling) observed in modern mid-ocean-ridge basalts (MORBs), but, in contrast to the pre-diction of the historical contingency model, these basalts do not exhibit suprasubduction zone–like geochemistry.

In the following sections, we present a more comprehensive review of the historical contingency model and its implications, and then we present detailed evidence that rebuts this proposal. We conclude that several processes can account for ophiolites, including formation at mid-ocean ridges, but that most large, structurally intact ophiolites with suprasubduction-zone geo-chemical signatures must have formed above active subduction zones. We further suggest that a more appropriate formulation of the “ophiolite conundrum” is this: Given that many, if not most, ophiolites have geochemical and petrologic signatures consis-tent with formation above active subduction zones, under what circumstances does this setting result in rock associations and structures consistent with those observed in ophiolites?

THE “OPHIOLITE CONUNDRUM” AND “HISTORICAL CONTINGENCY”

The “ophiolite conundrum” poised by Moores et al. (2000) addresses the structural and stratigraphic evidence in ophiolites that suggests formation by nearly 100% extension, which they suggest is uniquely characteristic of oceanic spreading ridges, and the overwhelming geochemical and petrologic evidence that these same rocks formed above subduction zones. Or as they state it: “This geochemical and petrologic evidence stands in strong contrast with evidence—from both the structure within Tethyan ophiolite complexes and the paleogeographic environ-ments inferred from surrounding and overlying sedimentary deposits—that these ophiolites originated well away from any type of subduction-related activity” (Moores et al., 2000, p. 4). As stated, the conundrum implies that the structural and sedi-mentary associations found in these ophiolites are inconsistent with any type of arc environment, and that this evidence is more compelling than the geochemical and petrologic evidence (see Pearce [2003] for a discussion Bayesian decision methods as applied to the ophiolite conundrum).

Ophiolites have long been used as natural laboratories for studying processes related to mid-ocean-ridge spreading and the generation of oceanic basalts. Thus, the ophiolite conundrum is an issue primarily to those who make direct correlations between the structural architecture of ophiolites and ocean crust formed at mid-ocean-ridge spreading centers. Nonetheless, the historical contingency hypothesis has implications beyond these correla-tions (which may, in any event, provide robust models despite the different origins inferred for ophiolites and ridges).

What is at stake in resolving the ophiolite conundrum? The heart of the issue is how we interpret the ophiolite record, in particular, with regard to paleotectonic reconstructions. Moores

LMLM

active island arc

active island arc remnant arcSSZ-like back arc spreading

SSZ forearc spreading

active island arc remnant arcMORB-like back arc spreading

AM

A

B

C

D

E

oceanic crust

AM

Figure 1. Western Pacifi c intra-oceanic subduction zone model of ophio lite formation by slab retreat and upper-plate extension (after Stern and Bloomer, 1992; Bloomer et al., 1995; Shervais, 2001). (A) Subduction initiates along transform margin, juxtaposing young, thin (hotter) oceanic lithosphere against older, thicker (cooler, more dense) oceanic lithosphere. (B) Initiation of a new subduction margin gener-ates suprasubduction-zone (SSZ) oceanic crust in a nascent forearc by extension and magmatism in response to slab sinking and hinge roll-back. (C) Stabilization of an island-arc crust by continued subduction. (D) Rifting and reconfi guration of the arc generates suprasubduction-zone oceanic crust by spreading in a narrow back-arc basin in response to continued hinge rollback. (E) Back-arc basin widens as hinge roll-back continues. Solid arrows are motion vectors of the subducting slab; dashed arrows denote migration of asthenosphere into space created by sinking slab and hinge rollback. In panels C–D, dashed outline denotes regions of hotter mantle beneath the regions of active arc and back-arc magmatism (after Weins and Smith, 2003). AM—astheospheric mantle, LM—lithospheric mantle, MORB—mid-ocean-ridge basalt.


et al. (2000) specifi cally addressed Tethyan ophiolites in devel-oping the historical contingency model; however, their model has gained attention not only with regard to Tethyan ophiolites (Barth et al., 2003; Liati et al., 2004; Beccaluva et al., 2004) but also with regard to ophiolites from a much broader temporal and geographic context (e.g., Moores, 2002, 2003; Saltus et al., 2003; Maxeiner et al., 2005).

The historical contingency model of Moores et al. (2000) consists of several disparate elements that must be discussed and understood separately. In its simplest proposition, Moores et al. state, “We propose a model wherein the nature of mantle tapped at mid-oceanic ridges has varied in the past in response to prior tectonic history of a region and/or the mantle” (Moores et al. 2000, p. 4). They suggest several mechanisms that may affect mantle source regions such that magmas formed at mid-ocean ridges, both current and ancient, carry the geochemical signature typical of modern subduction-zone magmas. These mechanisms include asthenosphere modifi ed by previous sub-duction events, formation of slab windows where ridges are subducted orthogonally or obliquely, the subducted slab com-ponent carried by old plates, and isotopically distinct compo-nents found in ocean-island basalts.

Suprasubduction Geochemistry in Basalts from Active Mid-Ocean Ridges

In support of their model, Moores et al. (2000) point to two modern plate confi gurations where mantle modifi ed by recent or active subduction could contaminate the source region of an active mid-ocean-ridge spreading center. One confi guration involves seafl oor spreading over a region of asthenospheric man-tle previously modifi ed by subduction fl uids, e.g., asthenosphere that recently resided in the mantle wedge region of a now extinct subduction zone. The clearest recent example of this is the Wood-lark basin in the southwest Pacifi c, where collision of the Ontong Java plateau with the Solomon arc stalled subduction along the NE margin of the arc and caused inception of a new, NE-dipping subduction zone along the SW margin of the arc (Taylor and Exon, 1987; Perfi t et al., 1987; Staudigel et al., 1987; Crook and Taylor, 1994; Johnson et al., 1987; Muenow et al., 1991). Thus, the Woodlark spreading center is a former back-arc basin that is now being subducted beneath the Solomon arc. This setting is complicated by the fact that the Woodlark spreading center is being subducted orthogonally, opening a slab window within the trench that could allow subduction components to migrate into the upper plate, as discussed subsequently.

A second confi guration is ridge-trench-trench triple junc-tions that mark the intersections of active mid-ocean-ridge spreading centers with active subduction zones. Modern exam-ples include the Woodlark (Perfi t et al., 1987), Chile (Klein and Karsten, 1995), and Juan de Fuca (Cousens et al., 1995) spread-ing ridges, and such confi gurations must have been common in the past (Klein and Karsten, 1995). Ridge subduction opens a slab window in the subducted plate because spreading contin-

ues prior to ridge subduction but stops once the ridge has been subducted. Formation of this slab window may allow communi-cation between subduction-modifi ed asthenosphere and astheno-sphere sources beneath the active spreading center prior to its subduction (Klein and Karsten, 1995; Cousens et al., 1995). This communication may result in magmas derived from an active spreading center that carry a subduction-like chemical and iso-topic compositional component (Klein and Karsten, 1995). Each of these ridges (Chile, Woodlark, and Juan de Fuca) was cited by Moores et al. (2000) as an example of modern mid-ocean ridges where suprasubduction-zone geochemical signatures have been detected. We address these examples and their signifi cance to the suprasubduction-zone ophiolite debate later in the paper.

The Fate of Old Plates: Subducted Slab Component and Long-Term Mantle Heterogeneities

Numerous studies have shown that the mantle sources for ocean island basalts (OIBs) are heterogeneous both in terms of trace-element enrichment and isotopic composition. Several distinct isotopic components have been proposed (depleted MORB mantle [DMM], high μ [HIMU], enriched mantle I [EMI], enriched mantle II [EMII], prevalent mantle [PREMA]) that must represent persistent, long-term trace-element hetero-geneities in the mantle (Zindler and Hart, 1986; Hart, 1988; Hofman, 1997). PREMA and DMM represent, respectively, the predominant isotopic composition of the mantle and the depleted, MORB-source asthenosphere. HIMU, EM1, and EM2 isotopic compositions record the recycling of altered oceanic lithosphere (including continent-derived sediments) and/or continental lithosphere (including subduction-modifi ed subcontinental lithospheric mantle) into the mantle either directly via subduction or potentially during entrainment of continental lithosphere during continental rifting.

Moores et al. (2000) presented numerical models of mantle dynamics suggesting that trace chemical and isotopic hetero-geneities related to subduction of oceanic lithosphere may per-sist on extremely long time scales. After Kellogg et al. (1999), they suggest that this subducted slab material accumulates in the mid-mantle region around 1500 km depth and that later it may be recycled and mixed into the overlying MORB-source asthenosphere, thus introducing a “subduction component” into normal mid-ocean-ridge basalts. Moores et al. (2000) further suggest that perturbations of the deeper mantle heterogeneities may be linked to cycles of continental assembly and dispersion (Wilson cycles), leading to periods during Earth history when increased mantle plume activity carries subduction-modifi ed mantle into the zone of MORB production.

An implication or unstated assumption of the historical contingency model is that when partial melting beneath mid-ocean-ridge centers taps mantle regions that carry isotopic evi-dence of subduction recycling, the resulting basalts should carry suprasubduction-zone trace-element signatures. This assumption is testable with geochemical and isotopic data from modern mid-


ocean-ridge basalts, much of it published over the last decade. In the following sections, we present data from modern ocean basins, including both mid-ocean ridges and suprasubduction zones, from well-studied examples of the ophiolite record, and from modern ridges with suprasubduction zone–like composi-tions, and we use those data to evaluate the historical contingency model and its impact on interpretations of the ophiolite record.

GEOCHEMISTRY OF MODERN MID-OCEAN-RIDGE AND SUPRASUBDUCTION-ZONE MAGMATIC ROCKS

An empirical relation between the composition of modern magma suites and tectonic setting was recognized soon after the emergence of the modern plate tectonic theory. Over the ensuing decades, much research has led to an understanding of the basic petrologic processes underlying these relationships. Magma-tism in the modern ocean basins can be considered to represent four major types: normal mid-ocean-ridge basalts (N-MORB), enriched mid-ocean-ridge basalts (E-MORB), within-plate or ocean-island basalts (OIB), and suprasubduction-zone magmas. Basaltic compositions dominate N-MORB, E-MORB and OIB suites; by contrast suprasubduction-zone suites are composition-ally more diverse, ranging from basalt to more evolved com-positions, and their plutonic equivalents, and they may include unusual lava compositions like boninites and adakites. Back-arc basin magmas form above subduction zones, but their compo-sitions are most similar to mid-ocean-ridge basalts when these basins are relatively mature.

Boninites are high-Si, high-Mg andesites that are found only in primitive or nascent arc terranes (e.g., Crawford and Falloon, 1989). Boninites are most commonly found, along with low-K tholeiites and felsic differentiates, in forearc terranes that repre-sent the extended basement upon which some modern arcs are built, e.g., the Izu-Bonin arc, the Marianas, and the Cape Vogel arc (Hickey and Frey, 1982; Bloomer and Hawkins, 1983; Walker and Cameron, 1983; Crawford and Falloon, 1989; Stern et al., 1991). These formed during rapid extension of the crust over a nascent subduction zone, prior to the establishment of modern arcs (Fig. 1B). Other boninites appear to form when a mantle plume or propagating back-arc basin rift extends into the forearc region of a modern arc (e.g., Deschamps and Lallemand, 2003). In all cases, experimental data suggest that boninites represent partial melts of highly depleted previously melted mantle, in response to high fl uid fl ux from the subducting slab (e.g., Falloon and Danyushevsky, 2000; Flower, 2003; Van der Laan et al., 1989; Umino and Kushiro, 1989). The high fl uid fl ux lowers the solidus of the refractory, enstatite-rich mantle; the resulting melts are rich in silica and MgO because enstatite (which is rich in both elements) dominates the melting assemblage. As a result, subarc lithosphere is dominated by a residual refractory mantle of harz-burgite composition, not lherzolite. In contrast, suboceanic litho-sphere is dominated by lherzolite (diopside-bearing peridotite), which represents smaller degrees of partial melting of the MORB asthenosphere source (Dick, 1989).

Major-Element Compositions of Mid-Ocean-Ridge and Suprasubduction-Zone Basalts

Oceanic basalts formed at normal or enriched mid-oceanic-ridge segments and ocean-island basalts formed within oceanic plates or on plume-enhanced ridge segments are characterized by a limited range in silica contents, moderate to high TiO

2 concen-

trations, and by tholeiitic fractionation trends. Table 1 compares the average compositions (along with minimum, maximum, and standard deviation) of over 2499 mid-ocean-ridge and 545 back-arc basin volcanic rocks from the PETDB database (Lehnert et al., 2000) to 1335 analyses from the Georoc database of basalts from primitive arcs that may be analogous to ophiolites (Mariana and Tonga arcs). The PETDB MORB database includes samples from slow (Indian ridge, Mid-Atlantic Ridge), intermediate (Juan de Fuca Ridge), and fast (East Pacifi c Rise) spreading centers. The contrasting compositions are highlighted in Figure 2, which compares the silica and TiO

2 concentrations of these suites.

A few lavas from the MORB database exhibit extreme compositions, but, in general, the data defi ne groups with con-sistent geochemical characteristics. Silica is uniformly low (48–52 wt% SiO

2), with little variation (mean ≈ 50 wt%; sigma ≈

1%); less than 2.5% of the data exceeds 52 wt% SiO2, and in no

case does silica exceed 65 wt% SiO2 (Table 1; Fig. 2A). Basalts

from mature back-arc basins have similar silica modes (Table 1; Fig. 2B). In contrast, volcanic rocks from primitive island arcs have a wide range in silica contents, with modes near 52 wt% SiO

2; almost half of the data has silica >52 wt% SiO

2 (Table 1;

Fig. 2C). Data for TiO2 present a similar picture: MORB values

range in TiO2 from ~0.5 wt% to 3.6 wt%, with a mode around

1.6 wt% TiO2 (Fig. 2D). Back-arc basin basalts have slightly

lower modes (≈1.3 wt% TiO2), while arc-volcanic rocks have

TiO2 modes around 0.9 wt% (Fig. 2E). More than 90% of arc-

volcanic rocks have TiO2 less than 1.4 wt%, whereas more than

80% of all ridge basalts have TiO2 greater than 1.2 wt% (Fig. 2F).

The major-element data reviewed here, along with those contained in the databases but not discussed here, show that mid-ocean-ridge basalts exhibit a restricted range in major-element concentrations, which are confi ned almost exclusively to compo-sitions that would be defi ned as tholeiitic basalts. More evolved compositions do occur, e.g., the recently described dacite samples on the Pacifi c-Antarctic ridge at 55–65 wt% SiO

2, which repre-

sent extreme fractionation of MORB parent magmas (Stoffers et al., 2002). However, these evolved compositions are rare and do not represent a signifi cant fraction of the magmatic activity at mid-ocean-ridge spreading centers. Thus, if the evolved magmas found in suprasubduction-zone ophiolites are to be interpreted as mid-ocean-ridge dacite, these magmatic rocks should represent only a small fraction of the total volcanic record, which must be dominated by lavas with normal or enriched MORB composi-tions. In contrast, if suprasubduction-zone ophiolites represent primitive-arc volcanism, evolved magmas should be more com-mon (up to half of all lavas), and the associated basalts should have arc tholeiite or calc-alkaline affi nities.


Trace-Element Signatures of Mid-Ocean Ridge and Suprasubduction-Zone Basalts

Potential problems with the stability of major-element concentrations during low-temperature hydrous metamorphism have long been recognized. As a result, most geochemical stud-ies of submarine volcanic rocks in ophiolites have focused on the trace composition of the basaltic rocks, which carry the most stable information regarding magma source and, by infer-ence, tectonic setting.

N-MORB, E-MORB, and OIBAverage N-MORB is often used as a reference when dis-

cussing the trace-element composition of oceanic magmas, including those from subduction settings. Compositional data for N-MORB rocks indicate a source depleted in incompatible trace elements relative to estimated primitive-mantle compositions. This source, referred to as depleted MORB mantle (DMM), is thought to reside in the shallow asthenosphere. Nd-Sr-Pb isotopic data suggest that DMM was formed during melt extraction events that occurred early in Earth history (Jacobsen and Wasserburg, 1980; O’Nions et al., 1977; Hofmann, 1997).

A standard method for looking at the trace-element compo-sition of oceanic basalts is the N-MORB–normalized spider dia-gram (Fig. 3) where trace elements are arranged from right to left in the order of increasing incompatibility with respect to mantle mineralogy. Average compositions for N-MORB, E-MORB, and OIB, taken from Sun and McDonough (1989), are shown on Fig-

ure 3A. Relative to N-MORB, OIB is systematically enriched in the more incompatible trace elements (smooth slope on left side of Fig. 3A). Ocean-island basalts are thought to be derived from a deeper, trace element–enriched mantle source, and they may refl ect a complex, multicomponent source that includes oceanic lithosphere recycled into the mantle by subduction. Average E-MORB is also enriched relative to N-MORB but to a lesser extent than OIB. The relative enrichments and depletions in oceanic basalts seen in the spider diagrams (Fig. 3A) can be conveniently illustrated using ratio-ratio plots such as the Nb/Yb versus Th/Yb plot (Pearce, 1982; Pearce et al., 1995) shown in Figure 3B. Such plots show the ratio of a more incompatible ele-ment (Nb, Th) and a less incompatible element (Yb); previous partial melting in a mantle source produces a decrease in both the Nb/Yb and Th/Yb ratios, while enrichments related to mantle plumes result in increases in both ratios. Thus, average N-MORB, E-MORB, and OIB magmas form a depletion-enrichment array on Nb/Yb versus Th/Yb plots.

Suprasubduction-Zone BasaltsThe hallmark of suprasubduction-zone basaltic magma

compositions is elevated concentrations of large ion litho-phile elements (LILE: Cs, Rb, Ba, Th, K, Sr, Pb) relative to high fi eld strength elements (HFSE: Nb, Ta, Hf, Zr, Ti) (Wood, 1980; Saunders et al., 1980; Pearce, 1982; Pearce et al., 1984). Subduction-zone models hold that fl uids and/or siliceous melts derived from the subducting oceanic slab carry high concen-trations of LILE (±light rare earth elements [LREEs]) that

TABLE 1. MEAN COMPOSITIONS, STANDARD DEVIATIONS, MAXIMUMS, AND MINIMUMS FOR MID-OCEAN-RIDGE BASALT (MORB) (MID-ATLANTIC RIDGE, EAST PACIFIC RISE, JUAN DE FUCA, INDIAN RIDGE), BACK-ARC BASINS, AND PRIMITIVE ARCS

(MARIANA, IZU-BONIN, VANUATU, TONGA, KERMADEC) Mid-Atlantic Ridge East Pacific Rise Juan de Fuca Indian Ridge Back-arc basins Mean Stddev Max Min Mean Stddev Max Min Mean Stddev Max Min Mean Stddev Max Min Mean Stddev Max Min SiO2 50.31 1.02 59.46 45.40 50.05 1.05 60.19 46.63 49.82 1.31 51.87 40.43 50.28 1.03 60.40 47.12 51.15 2.53 72.41 45.46 TiO2 1.47 0.34 2.86 0.45 1.76 0.41 3.32 0.68 1.66 0.43 3.53 0.97 1.61 0.51 3.82 0.52 1.30 0.43 2.77 0.35 Al2O3 15.43 0.95 22.10 10.80 14.75 1.30 20.43 0.13 14.96 1.16 17.53 10.69 15.70 1.25 21.30 12.68 15.98 1.29 25.53 7.23 FeOT 9.66 1.13 14.16 0.00 10.38 1.80 15.44 1.12 10.53 1.55 17.38 7.60 9.38 1.56 14.63 5.72 9.17 1.45 16.55 3.15 MnO 0.17 0.03 0.34 0.05 0.189 0.027 0.31 0.1 0.19 0.03 0.30 0.10 0.17 0.03 0.32 0.09 0.17 0.03 0.30 0.06 MgO 7.87 1.15 22.60 4.75 7.38 1.18 15.23 1.59 7.49 1.12 12.48 3.78 7.56 1.09 10.22 3.93 6.90 1.94 23.50 0.75 CaO 11.37 0.84 14.10 1.40 11. 37 0.90 13.35 4.48 11.61 0.79 13.32 8.38 11.00 0.84 13.84 7.76 10.94 1.51 13.86 3.03 Na2O 2.67 0.36 3.97 1.10 2.68 0.35 5.24 1.41 2.62 0.27 3.30 1.79 3.04 0.42 5.03 1.95 2.82 0.66 5.26 0.72 K2O 0.23 0.19 1.27 0.01 0.16 0.16 2.20 0.01 0.21 0.14 0.60 0.02 0.27 0.32 1.77 0.03 0.35 0.31 3.25 0.01 P2O5 0.16 0.06 0.42 0.03 0.17 0.07 1.15 0.04 0.17 0.08 0.59 0.05 0.20 0.09 0.67 0.04 0.17 0.10 1.12 0.01

n = 1396 n = 703 n = 106 n = 294 n = 545

Mariana Izu-Bonin Vanuatu Tonga Kermadec Mean Stddev Max Min Mean Stddev Max Min Mean Stddev Max Min Mean Stddev Max Min Mean Stddev Max Min SiO2 51.55 4.77 79.20 42.30 56.01 7.01 78.30 43.03 54.06 6.61 71.23 46.38 53.40 4.91 76.65 43.80 54.21 7.07 73.53 44.69 TiO2 0.99 0.37 2.55 0.17 1.08 0.68 2.70 0.05 0.83 0.38 2.49 0.44 1.11 0.61 2.51 0.14 0.83 0.35 1.68 0.12 Al2O3 15.30 1.87 20.18 10.18 15.02 1.81 20.57 11.43 15.54 2.17 19.66 8.84 15.61 2.34 25.20 0.16 16.31 2.33 20.57 5.50 FeOT 10.56 2.53 16.58 2.19 8.50 2.90 11.80 0.76 9.39 2.54 16.24 4.17 9.57 1.91 13.22 1.01 8.25 3.12 12.83 0.51 MnO 0.18 0.05 0.34 0.01 0.18 0.07 0.34 0.00 0.20 0.04 0.31 0.09 0.18 0.04 0.34 0.01 0.17 0.05 0.30 0.02 MgO 6.70 2.79 18.06 0.36 4.15 3.00 13.08 0.00 5.92 4.44 22.61 0.70 6.40 4.44 46.35 0.36 5.38 2.75 14.75 0.31 CaO 0.22 0.91 14.00 0.03 0.18 0.07 0.34 0.00 0.20 0.04 0.31 0.09 0.18 0.04 0.34 0.01 0.17 0.05 0.30 0.02 Na2O 2.52 0.65 4.72 0.41 3.17 1.19 6.90 0.39 2.84 0.81 5.20 1.43 2.61 1.03 6.39 0.10 2.48 1.08 4.93 0.00 K2O 0.62 0.53 3.82 0.00 1.16 1.36 11.51 0.20 1.89 1.15 4.90 0.19 0.42 0.36 1.91 0.01 0.52 0.51 2.69 0.01 P2O5 0.12 0.09 0.76 0.01 0.23 0.22 0.90 0.00 0.26 0.10 0.46 0.09 0.16 0.14 1.10 0.01 0.11 0.05 0.30 0.01

n = 245 n = 241 n = 566 n = 198 n = 85 Note: Back-arc basins include Mariana trough, North Fuji basin, Lau basin, Pearce Vela basin, West Philippine basin, Shikoku basin, Woodlark basin,

South Sandwich basin, Sulu Sea, and Bransfield Strait. FeOT —total Fe as FeO.


meta somatize the overlying mantle wedge and aid in lowering solidus temperatures. In addition to contributions from the slab itself (i.e., dehydration reactions in altered oceanic crust), sub-duction fl uids may carry elemental contributions from a variety of subducted sediments. HFSE and heavy (H) REE concentra-tions, on the other hand, are controlled primarily by the premeta-

somatism composition of the mantle wedge. Consequently, the trace-element signature of subduction-related magmas appears to be derived from three main sources: the subducted oceanic lithosphere, subducted sediment, and the mantle wedge over-lying the subducting slab (Perfi t et al., 1980; Pearce, 1982; Arculus and Powell, 1986; Davidson, 1996).

0

200

400

600

800

1000

1200

40 44 48 52 56 60 64 68 72 76

MORB SiO2

2469 analyses MAREPRIRJFR

Nu

mb

er o

f A

nal

yses

SiO2 (wt%)

0

10

20

30

40

50

40 44 48 52 56 60 64 68 72 76

Back Arc Basin SiO2

106 analyses BAB

Nu

mb

er o

f A

nal

yses

SiO2 (wt%)

0

50

100

150

200

250

40 44 48 52 56 60 64 68 72 76

Primitive Arcs SiO2

1657 analyses Marianas

Tonga

Nu

mb

er o

f A

nal

yses

SiO2 (wt%)

0

100

200

300

400

500

600

0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6 4.0 4.4 4.8

MORB TiO2

2469 analyses MAREPRIRJFR

Nu

mb

er o

f A

nal

yses

TiO2 (wt%)

0

20

40

60

80

100

120

140

0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6 4.0 4.4 4.8

Back Arc Basins TiO2

103 analysesBAB

Nu

mb

er o

f A

nal

yses

TiO2 (wt%)

0

50

100

150

200

250

300

350

400

0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6 4.0 4.4 4.8

Arc TiO2

1657 analyses Tonga

Marianas

Nu

mb

er o

f A

nal

yses

TiO2 (wt%)

A

B

C

D

E

F

Figure 2. Histograms illustrating SiO2 and TiO2 contents (wt%) in modern mid-ocean-ridge and subduction-zone basalts: (A) SiO2 in modern MORB, n = 2469; (B) SiO2 in modern back-arc basins, n = 106; (C) SiO2 in modern primitive arcs, n = 1657; (D) TiO2 in modern MORB, n = 2469; (E) TiO2 in modern back-arc basin basalts, n = 103; (F) TiO2 in modern primitive arcs, n = 1657. See Table 1 for complete major-element summary. BAB—Back Arc Basin, EPR—East Pacifi c Rise, IR—Indian Ridge, JFR—Juan de Fuca Ridge, MAR—mid-Atlantic Ridge.


Three samples from the South Sandwich Island arc are plotted (Pearce et al., 1995) in Figures 3C and 3D: a calc-alkaline basalt, a tholeiite basalt, and a low-K tholeiite basalt. Elevated concentrations of slab-derived components (Cs, Rb, Ba, Th, U, K, La, Ce, Pb, Sr) are superimposed on conservative mantle wedge-derived components (Nb, Zr, Sm, Eu, Ti, Dy, Y, Yb, Lu), pro-ducing the distinctive pattern of by “spikes” and “troughs” that is characteristic of suprasubduction-zone basalts (Fig. 3C). For the calc-alkaline basalt, the mantle wedge-derived components approximate N-MORB compositions, suggesting a mantle wedge with DMM composition (N-MORB normalized values for Nb, Nd-Lu ~ 1 in Fig. 3C). The tholeiite basalt and a low-K tholeiite samples, however, have conservative mantle wedge-derived components more depleted than N-MORB (Fig. 3C). A particu-lar feature of these basalts is that the most incompatible mantle-derived components (e.g., Nb) typically are more depleted than less incompatible mantle-derived components (e.g., Y, Yb). Such patterns cannot be produced by simple variations in the degree of partial melting and must represent a residual MORB mantle, i.e., one that has experienced a previous MORB melting event (Pearce and Parkinson, 1993). The Nb/Yb versus Th/Yb plot pro-vides a means for evaluating the mantle wedge contribution to the suprasubduction-zone basalt trace-element budget (Pearce, 1982). Addition of subduction-derived fl uid to the mantle wedge increases the Th/Yb ratio, but not the Nb/Yb ratio, producing the subduction component vectors shown in Figure 3D. Extrapola-tion back along a subduction component vector provides an indi-cation of presubduction composition of the mantle source.

Suggested tectonic settings for the generation of supra-subduction-zone ophiolites include back-arc basins, rifted island

arcs (the initiation of back-arc basin formation), and magmatic “forearcs” that form above nascent, retreating intraoceanic trenches (Fig. 1). In Figure 4, we have plotted forearc, arc, and back-arc trace-element data for magmatic rocks from three mod-ern oceanic subduction zones: Izu-Bonin-Mariana, New Britain–Manus, and Lau-Tonga. All of the samples in the data set have basalt or basaltic andesite compositions and classify as low-K tholeiites or boninites. The Izu-Bonin-Mariana data set includes Eocene-age forearc basalt and gabbro formed during initiation of subduction (Bloomer et al., 1995), Quaternary basalts from the active arc (Elliott et al., 1997), and Quaternary basalts from the active back-arc spreading ridge (Mariana Trough: Tian et al., 2005; Pearce et al., 2005). The New Britain–Manus data set (Woodhead et al., 1998) includes trench proximal (forearc) and arc basalts from New Britain and back-arc basalts of the active Manus spreading ridge, all Quaternary in age. The Lau-Tonga data set includes Miocene forearc basalt and gabbro formed during initiation of subduction, modern basalts from the active Tonga arc, and Pliocene to Quaternary basalts from the western Lau back-arc basin (Ewart et al., 1994).

Variations in the compositions of suprasubduction-zone basalts can be attributable to two effects: (1) differences in the magnitude of the subduction fl ux of LILEs and (2) verti-cal and/or lateral variations in the composition of the mantle wedge that control HFSEs (Taylor et al., 1992). The data in Figure 4 illustrate some of these variations. Although nearly all of the samples show signifi cant evidence of subduction-related LILE fl ux (prominent spikes and troughs in Figs. 4A, 4C, 4E, and 4G; elevated Th/Yb ratios in Figs. 4B, 4D, 4F, and 4H), some back-arc samples from each of the subduc-

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Pb Sr Nd Sm TiCs Ba U K Ce Pr P Zr Eu Dy Yb

Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

C D


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

B

OIB

EMORBNMORB

Calc-alkaline

Low-Ktholeiite

Tholeiite TholeiiteCalc-alkaline

Low-Ktholeiite

Figure 3. (A–B) Trace-element data for average normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) dis-played as N-MORB–normalized spider diagrams and Th/Yb-Nb/Yb ratio-ratio plots. (C–D) Trace-element data for typical subduction-zone basalts dis-played as N-MORB–normalized spider diagrams and Th/Yb-Nb/Yb ratio-ratio plots.


tion zones and three arc samples (two New Britain and one Tonga) plot on or near the mantle depletion-enrichment array on the Nb/Yb versus Th/Yb plot. Differences in the (presub-duction fl ux) composition of the mantle wedge are seen on the Nb/Yb versus Th/Yb plot as a broad range of Nb/Yb ratios and on the spider diagrams as a variable negative Nb anomaly and variable depletions in HFSEs (Nd through Lu) relative to N-MORB normalization.

In general, basalts erupted in a forearc setting, including boninites, are among the most depleted rocks on Earth. The forearc basalts from all three subduction zones are depleted in HFSEs relative to average N-MORB (NMORB–normal-ized values for Nb << 1.0, for Nd-Lu < 1.0, Fig. 4A; typically Nb/Yb < 0.8, Fig. 4B), which is consistent with derivation from a residual MORB mantle (RMM) source. The data for the arc basalts vary in HFSEs from depleted to enriched (Nb/Yb ≈ 0.15–1.1,


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Forearc

Arc

Forearc

Arc


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Izu - Bonin - Mariana

New Britain - Manus

Lau - Tonga

Backarc NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Backarc

BackarcBackarc

A B

C D

E F

G H

Figure 4. Trace-element variation for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) observed in: (A–B) Cenozoic forearcs, (C–D) Quater-nary arcs, (E–F) nascent back-arc basins, and (G–H) mature back-arc basins.


Figs. 4C and 4D). The Tonga arc basalts have uniformly depleted HFSE patterns, with Nb/Yb ratios below average N-MORB, consistent with a RMM source. Arc basalts from both the Mari-ana and New Britain systems show a range of HFSE patterns, from depleted to more enriched values, between those of aver-age N-MORB and E-MORB. With the exception of a few sam-ples from the New Britain arc, the arc basalts exhibit signifi cant subduction-related LILE enrichment. As a group, the back-arc basalts show the greatest variability both in the magnitude of LILE enrichment (Th/Yb ratios in Figs. 4B, 4D, 4F, and 4H) and the range of mantle enrichment-depletion in HFSEs (Nb/Yb ratio in Figs. 4B, 4D, 4F, and 4H). In general terms, back-arc basin basalt compositions vary with proximity to the spreading cen-

ter and the subducting slab; basalts erupted in proximal back-arc basins typically have arc-like trace-element patterns with an RMM or DMM mantle component, while more distal (wider) back-arc basins show progressively less subduction infl uence and more N-MORB-like and E-MORB-like trace-element composi-tions (cf. Figs. 4E and 4F with Figs. 4G and 4H).

OPHIOLITE GEOCHEMISTRY

We have compiled whole-rock geochemical data from eight well-studied ophiolites to illustrate the range of compo-sitions expressed in the ophiolite record. Concentrations of the major elements are shown in Figure 5 and summarized in

0

5

10

15

20

25

40 42 44 46 48 50 52 54 56 58 60 62 64 66 68 70 72 74 76 78

SSZ Ophiolite SiO2

Trinity

Betts Cove

Troodos

Num

ber

of A

naly

ses

SiO2

n = 195

0

10

20

30

40

50

60

70

0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6

SSZ Ophiolites TiO2

Trinity

Betts Cove

Troodos

Num

ber

of A

naly

ses

TiO2

n = 195

0

5

10

15

20

25

30

40 42 44 46 48 50 52 54 56 58 60 62 64 66 68 70 72 74 76 78

Mixed Ophiolite SiO2

JosephineBay of IslandsOman

Num

ber

of A

naly

ses

SiO2

n = 185

0

5

10

15

20

25

30

35

40

0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6

Mixed Ophiolites TiO2

Josephine

Bay of Islands

Oman

Num

ber

of A

naly

ses

TiO2

n = 185

0

1

2

3

4

5

6

7

8

40 42 44 46 48 50 52 54 56 58 60 62 64 66 68 70 72 74 76 78

“MORB” Ophiolite SiO2

Macquarie Island Pindos

Num

ber

of A

naly

ses

SiO2

n = 29

0

1

2

3

4

5

6

7

8

0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6

“MORB” Ophiolite TiO2

Macquarie Island

Pindos

Num

ber

of A

naly

ses

TiO2

n = 29

A

C

E

B

D

F

Figure 5. Histograms illustrating SiO2 and TiO2 contents (wt%) in ophiolite basalts: (A–B) SiO2 and TiO2 in true suprasubduction-zone (SSZ) ophiolites, n = 195; (C–D) SiO2 and TiO2 in mixed suprasubduction zone–mid-ocean-ridge basalt (MORB) ophiolites, n = 185; (E–F) SiO2 and TiO2 in MORB-only ophiolites, n = 29. See Table 2 for com-plete summary of major elements in these ophiolites.


Table 2, which reports their mean, minimum, maximum, and standard deviation. Trace-element geochemical data are plotted in Figures 6, 7 and 8; geologic and geochemical characteristics of the ophiolites are summarized in Table 3.

Tethyan ophiolites are represented by the Cretaceous Troodos ophiolite (Robinson et al., 1983; Robinson and Malpas, 1990; Malpas and Langdon, 1984; Cameron, 1985; Rautenschlein et al., 1985; Laurent, 1992; Baragar et al., 1990), the Cretaceous Oman ophiolite (Glennie et al., 1974; Alabaster et al., 1982; Lippard et al., 1986; Ernewein et al., 1988; Rochette et al., 1991; Nehlig, 1993; Ishikawa et al., 2002), and the Triassic Argolis (Pindos basin) ophiolite (Saccani et al., 2003). Cor dillera ophio-lites are represented by the Silurian-Devonian Trinity ophiolite (Wallin and Metcalf, 1998; Metcalf et al., 2000) and the Jurassic Josephine ophiolite (Harper, 2003a, 2003b). The Ordovician Bay of Islands ophiolite (Jenner et al., 1991) and the Ordovician Betts Cove ophiolite (Bedard et al., 1998; Bedard, 1999) repre-sent Appalachian (Iapetus) ophiolites. The Miocene Macquarie

Island ophiolite (Kamenetsky et al., 2000; Varne et al., 2000) is from the southwest Pacifi c Ocean. The data set includes examples of ophiolites with clear suprasubduction zone com-positions (Trinity, Betts Cove, Troodos), ophiolites with mixed suprasubduction-zone and MORB compositions (Josephine, Oman, Bay of Islands), and ophiolites with MORB compositions (Argolis, Macquarie Island).

Suprasubduction-Zone Compositions in Ophiolites

The Trinity, Betts Cove, and Troodos ophiolites all show similar suprasubduction-zone compositions. All three ophio-lites exhibit a broad range of silica compositions dominated by basalt to basaltic andesite, but they also include more silica-rich dacite and rhyolite compositions. The basaltic rocks are low in titanium—typically less than 1.2 wt% TiO

2—and are classifi ed

as low-K tholeiites to boninites (Fig. 5; Table 2). In general, the HFSEs are more depleted than N-MORB, with Nb/Yb ratios at

TABLE 2. MEAN COMPOSITIONS, STANDARD DEVIATIONS, MAXIMUMS, AND MINIMUMS FOR SUPRASUBDUCTION-ZONE (SSZ) OPHIOLITES (TRINITY, BETTS COVE, TROODOS), MIXED OPHIOLITES

(JOSEPHINE, OMAN, BAY OF ISLANDS), AND MID-OCEAN-RIDGE BASALT (MORB) OPHIOLITES (ARGOLIS, MACQUARIE ISLAND)

sodoorTevoCstteBytinirT Mean Std-dev Max Min Mean Std-dev Max Min Mean Std-dev Max Min SiO2 58.32 10.27 78.35 46.8 54.09 6.10 79.76 43.42 56.47 3.21 66.61 51.34 TiO2 0.428 0.2476 1.273 0.063 1.10 0.79 2.94 0.07 0.31 0.41 1.93 0.01 Al2O3 16.38 1.9418 10.98 19.54 15.50 2.04 20.96 8.91 1.06 0.28 1.51 0.45 FeO* 5.282 2.3265 0 0 9.35 2.57 20.58 0.75 15.70 0.90 18.21 13.45 MnO 0.111 0.0783 0.6 0. 013 0.18 0.09 1.16 0.03 10.02 1.78 14.87 5.94 MgO 6.297 4.3338 15.26 0.023 7.67 3.71 26.70 0.22 0.16 0.06 0.37 0.06 CaO 8.828 4.0171 18.05 1.463 7.70 3.19 16.16 0.08 5.37 1.52 9.06 1.57 Na2O 2.666 1.6295 6.443 0.011 3.64 1.50 7.31 0.04 6.82 2.70 17.39 1.77 K2O 0.126 0.1797 1.112 0.01 0.68 1.04 5.40 0.00 4.01 1.42 6.91 1.43 P2O5 0.10 0.13 1.02 0.00 0.08 0.03 0.16 0.03

n = 63 n = 214 n = 96

Josephine Oman V1 (Geotimes) Oman V2 (Lasail-Alley) Mean Std-dev Max Min Mean Std-dev Max Min Mean Std-dev Max Min SiO2 53.15 4.45 66.61 41.87 54.66 4.58 69.21 45.97 58.02 9.58 81.57 40.84 TiO2 1.26 0.85 3.46 0.23 1.46 0.39 2.44 0.51 0.65 0.26 1.28 0.20 Al2O3 14.80 1.72 18.03 10.22 15.47 1.39 18.49 11.64 14.73 2.09 19.32 8.41 FeO* 9.73 2.93 17.71 2.67 9.46 1.47 12.83 4.94 7.37 2.21 12.36 1.57 MnO 0.19 0.07 0.35 0.04 0.19 0.07 0.50 0.10 0.15 0.07 0.47 0.03 MgO 8.05 3.79 18.75 1.61 3.98 1.50 7.70 0.70 5.21 3.19 15.29 0.14 CaO 7.86 2.35 13.16 3.09 6.97 3.32 20.95 1.91 8.29 5.08 24.36 0.69 Na2O 3.90 1.47 8.33 0.99 5.48 1.23 8.31 0.22 3.63 1.63 7.37 0.0 K2O 0.62 0.65 2.99 0.01 0.30 0.36 1.88 0.01 0.50 0.96 7.91 0.0 P2O5 0.13 0.07 0.34 0.02 0.19 0.07 0.37 0.04 0.10 0.14 1.10 0.0

n = 49 n = 102 n = 76

dnalsIeirauqcaMsilogrAsdnalsIfoyaB Mean Std-dev Max Min Mean Std-dev Max Min Mean Std-dev Max Min SiO2 61.56 10.99 79.90 43.83 48.77 3.43 58.55 43.53 49.34 0.73 51.14 47.37 TiO2 0.98 0.83 3.44 0.12 1.60 0.77 2.89 0.21 1.51 0.22 2.10 0.97 Al2O3 15.83 2.69 19.75 11.07 15.05 0.87 16.55 12.82 16.98 0.69 18.22 15.03 FeO* 5.73 3.76 14.47 0.58 10.22 1.84 13.58 7.76 7.92 0.67 10.17 6.81 MnO 0.12 0.11 0.60 0.01 0.23 0.13 0.48 0.10 7.34 0.69 8.75 5.65 MgO 3.49 3.03 8.48 0.21 6.76 1.11 10.43 4.09 0.13 0.03 0.18 0.07 CaO 5.30 3.75 13.77 0.32 10.23 1.68 13.12 4.90 11.43 0.92 13.53 9.81 Na2O 5.59 1.79 9.48 2.75 3.61 0.86 4.98 1.08 3.11 0.43 4.24 2.37 K2O 1.19 2.19 11.01 0.01 0.39 0.41 1.39 0.04 0.74 0.34 1.76 0.12 P2O5 0.32 0.12 0.66 0.08

n = 40 n = 23 n = 55 Note: Data are from: Metcalf et al. (2000), Bedard (1999), Robinson et al. (1983), Robinson and Malpas (1990),

Malpas and Langdon (1984), Cameron (1985), Rautenschlein et al. (1985), Laurent (1992), Baragar et al. (1990), Harper (2003a, 2003b), Alabaster et al. (1982), Lippard et al. (1986), Einaudi et al. (2000), Jenner et al. (1991), Saccani et al. (2003), Kamenetsky et al. (2000). FeO*—total Fe as FeO.


or less than N-MORB values, while LILE/HFSE ratios are ele-vated (Figs. 6A–6F). Data from these three ophiolites are most similar to modern forearcs and suggest melting of a residual MORB mantle enriched by subduction fl uids. Geologic con-straints for both the Betts Cove and Trinity ophiolites are consis-tent with generation in a suprasubduction-zone forearc setting (Bedard, 1999; Metcalf et al., 2000). Generation of the Troodos ophiolite has been ascribed to various suprasubduction-zone

settings, including a forearc, a back-arc, or a marginal basin similar to the modern Andaman Sea (McCulloch and Cameron, 1983; Gass et al., 1984; Moores et al., 1984). Indeed, it was the major-element geochemistry of Troodos volcanic rocks that led Miyashiro (1973) to challenge the mid-ocean-ridge origin of Troodos (see Cann [2003] and Robinson et al. [2003], for discussions of the importance of Troodos in the development of the ophiolite concept).


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Betts Cove ophioliteBetts Cove ophiolite

Troodos ophioliteTroodos ophiolite

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Bay of Islands ophioliteBay of Islands ophiolite

C D

E F

G H


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

Trinity ophioliteA

Nb/Yb

NMORB

EMORB

OIB

0.01 0.1 1 10 100

10

1

0.1

0.01

Th/Yb

Trinity ophioliteB

Figure 6. Trace-element data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) from ophiolites with supra-subduction-zone trace-element chemis-try: (A–B) Trinity, (C–D) Betts Cove, (E–F) Troodos; and (G–H) a mixed suprasubduction zone–mid-ocean-ridge basalt (MORB) ophiolite: Bay of Islands.


Mixed Suprasubduction Zone and Mid-Ocean Ridge Compositions in Ophiolites

The Josephine, Oman, and Bay of Islands ophiolites are all examples of ophiolites that exhibit both suprasubduction-zone and MORB geochemical signatures (Fig. 5; Table 2). In the Josephine ophiolite, suprasubduction zone–like lavas are overlain by MORB-like lavas in the volcanic stratigraphy (Figs. 7A–7D). Basaltic

rocks in the lower lava sequence have low TiO2 values (typically

< 1.2 wt%), HFSE concentrations generally less than average N-MORB, and elevated LILE/HFSE ratios. Basaltic rocks of the upper lava sequence have higher TiO

2 values (typically > 1.2 wt%)

and HFSE concentrations and LILE/HFSE ratios that approximate N-MORB. Overall SiO

2 values of both the lower and upper lavas

range between 47 wt% and 53 wt%, with a few samples at higher values (up to 62 wt%). The Josephine ophiolite occupies a paleo-


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

Josephine ophioliteLower lavas &intrusions

Josephine ophioliteUpper lavas


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

Oman ophioliteLower lavas

Oman ophioliteUpper lavas

A

C

E

G

Josephine ophioliteLower lavas &intrusions

Josephine ophioliteUpper lavas

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Oman ophioliteLower lavas

Oman ophioliteUpper lavas

B

D

F

H

Figure 7. Trace-element data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) from ophiolites with mixed mid-ocean-ridge and suprasubduction-zone trace-element chemistry: (A–D) Josephine upper and lower lavas and (E–H) Oman upper and lower lavas.


position between two segments of a rifted Jurassic arc system and records spreading in an extensional back-arc setting where supra-subduction zone–like magmas give way to MORB-like magmas (Harper, 2003a, 2003b), similar to that seen in the modern Lau Basin and Mariana Trough (Hawkins, 2003; Pearce et al., 2005).

In the Oman ophiolite, MORB-like lavas (Figs. 7E–7F) are overlain by suprasubduction zone–like lavas (Figs. 7G–7H) in the volcanic stratigraphy. Basaltic to andesitic rocks of the lower lava have high TiO

2 values (typically > 1.2 wt%), and

HFSE concentrations and LILE/HFSE ratios that are similar to

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

Pindos basin Pindos basin

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Macquarie Island ophiolite

A B

D


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

Macquarie Island ophioliteC

Figure 8. Trace-element data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) from ophiolites with mid-ocean-ridge trace-element chemistry: (A–B) Pindos and (C–D) Macquarie Island (actually an uplifted segment of ridge, not a true ophiolite).

TABLE 3. SUMMARY OF MAIN CHARACTERISTICS OF EIGHT OPHIOLITES DISCUSSED IN TEXT Ophiolite age (Ma)

Locality Reported range of SiO2

(wt%)

Basalt geochemistry

Trace-element signature Mantle source

Frac. sequence

Cover sequence Paleotectonic interpretation

Josephine(162–164)

Cordillera 46–58 Lower: Low-Ti thol.

BoniniteUpper:

High-Ti thol.

Lower: High LILE/HFSE HFSE < N-MORB

Upper:LILE/HFSE ~ N-MORB

HFSE ~ N-MORB

Lower: RMM

Upper:DMM

B1 Volcaniclastic SSZ back arc

Trinity (431–398)

Cordillera 46–57 71–78

Low-Ti thol. High LILE/HFSE HFSE < N-MORB

RMM B1 Volcaniclastic SSZ forearc

Betts Cove (489)

Appalachia 46–59 Low-Ti thol. Boninite

High LILE/HFSE HFSE < N-MORB

RMM B1 Volcaniclastic SSZ forearc

Bay of Islands (484)

Appalachia 48–55 60–6472–78

Low-Ti thol. High-Ti thol.

Elevated LILE/HFSE HFSE ~ N-MORB

FMM to EMM A1, B1 Clastic Mature back arcor

Mid-ocean ridgeTroodos(92–90)

Tethys 49–65 Low-Ti thol. Boninite

High LILE/HFSE HFSE < N-MORB to ~ N-MORB

RMM to FMM B1,2 Chert overlain by marine carbonate

SSZforearc

Oman(97–94)

Tethys 45–77 Lower: High-Ti thol.

Upper:Low-Ti thol.

Lower: Elevated LILE/HFSE

HFSE ~ N-MORBUpper:

High LILE/HFSE HFSE < N-MORB

DMM to RMM Primary: B1,2

Minor:A1

Ophiolite breccia, marine carbonate

Mid-ocean ridgeand

SSZ forearc

Argolis(Triassic)

Tethys High-Ti thol. LILE and HFSE ~ N-MORB to E-MORB

DMM to EMM A Radiolarian cherts Mid-ocean ridge

Macquarie(9)

Pacific 47–51 High-Ti thol. LILE ~ E-MORB to OIB HFSE ~ E-MORB to OIB

EMM to OIB A1,2 Volcaniclastic w/ ophiolite clasts;

marine carbonate

Mid-ocean ridge

Note: MORB—mid-ocean-ridge basalt; LILE—large ion lithophile element; HFSE—high field strength element; N—normal; E—enriched; OIB—ocean-island basalt. Mantle source: DMM—depleted MORB mantle; RMM—residual MORB mantle; EMM—enriched MORB mantle. Fractionation sequence: A—olivine > plagioclase > clinopyroxene; B—olivine > clino/orthopyroxene > plagioclase > hornblende. 1—Based on cumulate sequences; 2—based on phenocryst assemblages.


N-MORB. Highly variable K contents in the lower lavas may result from hydrothermal alteration. Basaltic to andesitic rocks of the upper lava sequence have lower TiO

2 values (typically

< 1.2 wt%), HFSE concentrations that are generally less than aver-age N-MORB, and elevated LILE/HFSE ratios. While a majority of volcanic and plutonic rocks in the Oman ophiolite have SiO

2

values between 47 wt% and 58 wt%, a number of samples range between 60 wt% and 78 wt% SiO

2 (Fig. 5). The paleotectonic

setting of the lower lavas in the Oman ophiolite remains contro-versial (e.g., Ernewein et al., 1988; Shervais, 2001).

Although the Bay of Islands ophiolite exhibits both supra-subduction zone–like and MORB-like geochemical features (Figs. 6G–6H), there is no clear stratigraphic sequence to the magma types (Jenner et al., 1991). Overall, the ophiolite records a broad range of silica values with three modes, 48–55, 60–64, and 72–78 wt% SiO

2 (Table 3). Unlike the Josephine

and Oman ophiolites, which show trace-element compositions defi ning two distinct magma types, basaltic rocks of the Bay of Islands ophiolite exhibit a continuum of compositions between suprasubduction-zone and MORB types. For example, TiO

2 val-

ues span a broad range from low-Ti, suprasubduction zone–like values (<1.2 wt%) to high-Ti MORB-like values (1.2–2.2 wt%; Tables 2 and 3). A majority of the basalt samples show vary-ing degrees of subduction enrichment (variably elevated LILE/HFSE ratios, Figs. 6G–6H) superimposed on a mantle-derived component (HFSE) that varies between N-MORB and E-MORB. Another subset of basalt samples shows little or no subduction enrichment (Figs. 6G–6H), and several samples approximate E-MORB and OIB compositions. Generation of the Bay of Islands ophiolite has been ascribed to a back-arc basin (Jenner et al., 1991).

Mid-Ocean-Ridge Compositions in Ophiolites

Ophiolites exhibiting strictly MORB geochemical sig-natures are rare and are known largely from thrust slices and mélange blocks in accretionary complexes. Tethyan ophiolites from the Eastern Mediterranean region provide the best record of MORB compositions in the ophiolite record. Saccani et al. (2003) reported MORB-like compositions from the Triassic Argolis ophiolite in Greece (Figs. 8A–8B). The Argolis ophio-lite is one of several ophiolitic massifs that provide a record of the Mesozoic Pindos ocean basin, which formed during rift-ing along the northern margin of Gondwanaland. The Triassic Argolis ophiolite massif provides a record of the early develop-ment of the Pindos basin. The Argolis ophiolite is dominated by basaltic compositions (45–51 wt% SiO

2, with a few samples

at ~60 wt%) and TiO2 contents at 1.4–2.8 wt% (Saccani et al.,

2003). Trace-element ratios Th/Y-Nb/Y plot along the mantle depletion-enrichment array between N-MORB and E-MORB, albeit with slightly elevated Th/Yb ratios (Figs. 8A–8B). Trace-element patterns on the spider diagram are transitional between N-MORB and E-MORB; element mobility is apparent in a few elements (e.g., K, U).

Macquarie Island is a slice of Miocene ocean fl oor that lies above sea level in the Southern Ocean south of Tasmania (Varne et al., 2000). It can be argued that Macquarie Island is not, strictly speaking, an ophiolite because it has not yet been emplaced into continental or arc crust. Some workers, however, have regarded Macquarie Island as an example of a MORB-type ophiolite (e.g., Dilek, 2003); we include it here as such. Ocean fl oor exposed on Macquarie Island was generated by spreading on the Australia-Pacifi c spreading ridge at 9–14 Ma, and it was exposed during transpression that formed the Macquarie Ridge in the last 10 m.y. (Varne et al., 2000; Kamenetsky et al., 2000). The Macquarie Island ophiolite is dominated by basaltic com-positions (48–51 wt% SiO

2) with TiO

2 contents of 1–2 wt%.

Basalts from Macquarie Island are highly enriched in incom-patible trace elements, ranging from compositions equivalent to average E-MORB to compositions more enriched than OIB, and they plot along the depletion-enrichment array on the ratio-ratio plot (Fig. 8D). Trace-element patterns on the spider dia-grams closely parallel those of average E-MORB and average OIB (Fig. 8C).

ANOMALOUS “SUPRASUBDUCTION ZONE–LIKE” BASALTS AT ERUPTED MID-OCEAN RIDGES

A major thrust of the historical contingency model focuses on examples of modern mid-ocean ridges where “suprasub-duction zone–like” trace-element signatures have been reported. Each of these examples represents a ridge-trench-trench triple junction where an active spreading ridge intersects an active sub-duction zone, potentially permitting communication of MORB and suprasubduction-zone mantle source regions via a slab window. Three such examples were discussed by Moores et al. (2000): the Chile Ridge (Klein and Karsten, 1995; Karsten et al., 1996; Sturm et al., 1999, 2000), the Juan de Fuca Ridge (Cousens et al., 1995), and the Woodlark spreading center (Perfi t et al., 1987; Staudigel et al., 1987).

Chile Ridge

Klein and Karsten (1995) fi rst reported suprasubduction zone–like trace-element signatures for basalts collected from four active segments of the Chile Ridge adjacent to the ridge-trench-trench with the Andean subduction zone. Subsequent papers (Karsten et al., 1996; Sturm et al., 1999, 2000) have reported Sr-Nd-Pb isotopic data and discussed the occurrence in the con-text of the ophiolite conundrum. Trace-element data for the Chile Ridge are shown in Figures 9A and 9B. Enrichments of LILEs (Cs, Rb, Ba, Th, K, Pb, and Sr) are evident on both the Th/Y-Nb/Y ratio plot and the spider diagrams, superimposed on mantle com-ponents that vary from slightly depleted N-MORB to E-MORB. A few samples show, relative to average N-MORB, a weak Nb negative anomaly and a slight enrichment of LILE. Samples at the other end of the spectrum show, relative to E-MORB, slight enrichments in LILEs that also produce a weak negative Nb


anomaly. Although present, the suprasubduction-zone signature in Chile Ridge basalts is subtle when compared to that of either modern subduction zones or the suprasubduction-zone ophiolite record (cf. Figs. 9A and 9B to Figs. 6, 7, and 8).

Juan de Fuca Ridge

Cousens et al. (1995) reported geochemical and isotopic data from the West Valley segment of the Juan de Fuca Ridge near the ridge-trench-trench junction with the North American plate. Trace-element patterns on the spider diagrams parallel those of E-MORB and show no enrichment in LILEs relative to HFSEs (Fig. 9C). On the Th/Y-Nb/Y ratio plot, the Juan de Fuca data plot along the enrichment-depletion centered on E-MORB (Fig. 9D). Cousens et al. (1995) used Nd-Sr-Pb isotopic compositions to evaluate mantle source regions and found evidence of a hetero-geneous source formed by mixing of DMM and HIMU compo-nents. This locality provides no evidence for a suprasubduction

zone–like component at an active mid-ocean-ridge spreading center because there is very little about the composition of these basalts that are suprasubduction zone–like. The Juan de Fuca data do provide additional evidence that mantle source regions modi-fi ed by subduction recycling do not necessarily give rise to basalts with suprasubduction zone–like trace-element compositions.

Woodlark Basin

Perfi t et al. (1987) and Staudigel et al. (1987) reported geochemical data for recent basalts collected from the active Woodlark basin spreading ridge. In addition to its position at a ridge-trench-trench triple junction, recent subduction reversal places the Woodlark spreading ridge over mantle that has for-merly been modifi ed by subduction processes (Perfi t et al., 1987; Staudigel et al., 1987). Woodlark basalts have Nb/Yb ratios simi-lar to N-MORB, and some samples show slightly elevated Th/Yb ratios; trace-element patterns on the spider diagram approximate


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

Woodlark basin

Chile RidgeChile Ridge

Juan de Fuca RidgeJuan de Fuca Ridge

Woodlark basin

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A B

C D

E F

Figure 9. Trace-element data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) for basaltic rocks collected from modern mid-ocean ridges with re-ported suprasubduction zone–like com-positions: (A–B) Chile Ridge, (C–D) Juan de Fuca Ridge, and (F–G) Wood-lark basin.


N-MORB (Figs. 9E–9F). The Woodlark basalts carry a weak sub-duction signature similar to that seen in mature back-arc basins (cf. Figs. 9E–9F with Figs. 4E–4H).

MANTLE HETEROGENEITY: TRACE-ELEMENT AND ISOTOPIC VARIATIONS IN MODERN MID-OCEAN-RIDGE BASALTS

A main thesis of the historical contingency model is that mantle source regions exhibit long-lived heterogeneities related to prior history, specifi cally the recycling of oceanic lithosphere into the mantle via subduction. The model further suggests that while modern mid-ocean ridges primarily tap DMM mantle sources, past mid-ocean ridges may have tapped subduction-modifi ed mantle heterogeneities, thus making geochemical data unreliable as an indicator of tectonic setting and by infer-ence ophiolite discrimination. Early recognition of subduction-modifi ed mantle heterogeneities was found largely in iso topic data from ocean-island and plume related basalts (Zindler and Hart, 1986; Hart, 1988) as noted by Moores et al. (2000). An unstated assumption of the historical contingency model is that partial melting of mantle regions carrying isotopic evi-dence of subduction recycling would produce basalts that carry suprasubduction-zone trace-element signatures.

It is widely recognized that magmas erupted at modern mid-ocean ridges are quite variable in terms of both trace-element and isotopic composition. These variations largely refl ect source hetero geneities and are observed at both regional (ocean basin) and local (single ridge segment or adjacent ridge segments)

scales. Isotopic data point to subduction recycling as a major contributor to MORB source heterogeneity, including most of the isotopic components identifi ed in the historical contingency model (e.g., HIMU, EMI, EMII). Thus, a test of the historical contingency model can be found in trace-element data from modern mid-ocean-ridge basalts that carry isotopic evidence of subduction contamination of their mantle source regions. In the following sections, we review data from several recent trace-element and isotopic studies of mid-ocean-ridge basalts. Despite trace-element and isotopic evidence for a heterogeneous source related to recycled (subducted) oceanic lithosphere, these basalts bear little resemblance to basalts from either modern subduction zones or the suprasubduction-zone ophiolite record.

North Chile Ridge

Recent basalts from the North Chile ridge (latitude 37–39°S; Bach et al., 1996) provide an example of mid-ocean-ridge basalts that have more depleted compositions than average N-MORB (Fig. 10). These data show a depletion in the most incompatible trace elements (Rb to Nb; Fig. 10A) and, in particular, a depletion in Nb relative to other HFSEs (e.g., Zr, Hf, Y, Yb). The depleted nature of these basalts relative to average N-MORB is particu-larly apparent in the Th/Yb-Nb/Yb ratio plot (Fig. 10B). Bach et al. (1996) attributed this depletion to prior removal of a low melt fraction from the N-MORB source mantle within the last few million years. Dynamic melting, i.e., episodic melt extrac-tion from a common source during a single protracted melting event, is capable of producing magmas from a common source

North Chile Ridge

Th/Yb

17.0 18.0 19.0 20.0 21.0

15.4

15.5

15.6

15.7

15.8

HIMU

EMII

EMI

DMM

207 P

b/20

4 Pb

143 N

d/14

4 Nd

206Pb/204Pb 87Sr/86Sr

C

0.702 0.703 0.704 0.705 0.706

0.5124

0.5126

0.5128

0.5130

0.5132

0.5134 DMM

HIMU

EMI

EMII

D


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

B

Figure 10. Trace-element and isotopic data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) for ba-saltic rocks from the North Chile mid-ocean ridge. DMM—depleted MORB mantle, HIMU—high μ, EMI—enriched man tle I, EMII—enriched mantle II.


that exhibit varying degrees of depletion (Pearce et al., 1995). Nd-Sr-Pb isotopic data (Bach et al., 1996) suggest mixing of DMM, EMII, and possibly HIMU mantle reservoirs (Figs. 10C–10D).

East Pacifi c Rise

Recent basalts from the northern East Pacifi c Rise (latitude 10–11°N; Niu et al., 1999; Regelous et al., 1999) exhibit a com-plete range of compositions between N-MORB and E-MORB (Figs. 11A–11B). On the ratio-ratio plot, data plot clearly along the mantle depletion-enrichment array, where the majority of samples are slightly more enriched than N-MORB (Fig. 11B). Correlations among trace-element ratios and Nd-Sr-Pb isotopic ratios point to mixing of two mantle components, one more trace element–depleted and one more trace element–enriched (Niu et al., 1999). Both mantle components are linked to recycled oceanic lithosphere—the depleted source to subducted litho-spheric mantle and the enriched component to subducted oceanic crust (Niu et al., 1999). The Nd-Sr-Pb isotopic compositions of basalts along this section of the East Pacifi c Rise represent a mixture of DMM and EMII mantle reservoirs (Niu et al., 1999), although component HIMU mantle cannot be ruled out.

South Atlantic Ridge

The South Atlantic Ridge was formed by rifting of continental lithosphere. Trace-element data (le Roux et al., 2002) for a suite of basalts collected from several segments of the active South Atlantic Ridge (latitude 40–52.5°S) show considerable variation but can be

broadly divided into two groups (Figs. 12A–12B), an N-MORB group (fi lled squares) and a group that trends toward more enriched compositions, overlapping E-MORB (open squares). Although both groups show slightly elevated Th/Yb ratios, more pronounced in the N-MORB group, the data generally plot along the mantle depletion-enrichment array on the Th/Y-Nb/Y ratio plot and exhibit clear N-MORB and E-MORB patterns on the spider dia-grams. Nd-Sr-Pb isotope data for the South Atlantic Ridge basalts provide evidence of a heterogeneous mantle source (le Roux et al., 2002). These data suggest mixing of DMM, EMI, EMII, and/or HIMU isotopic components in the mantle source the South Atlantic Ridge (Figs. 12C and 12D). le Roux et al. (2002) interpreted these iso topic signatures to refl ect the infl uence of (1) altered oceanic lithosphere and pelagic sediment recycled to the shallow mantle via plumes, (2) remnants of delaminated subcontinental litho-spheric mantle, and (3) Mesozoic suprasubduction metasomatism of mantle beneath Gondwanaland prior to opening of the South Atlantic. Trace-element variations discussed previously correlate closely with the observed isotopic variations, providing a link between trace-element compositions and mantle heterogeneities related to subduction recycling (le Roux et al., 2002).

Australian-Antarctica Discordance

A section of the Southeast Indian Ridge between the rifted margins of Australia and Antarctica marks the boundary between Indian-type and Pacifi c-type mantle domains and has been referred as the Australian-Antarctica discordance. Pyle et al. (1992) used isotopic data to map the location of the Australian-Antarctica dis-

17.0 18.0 19.0 20.0 21.0

15.4

15.5

15.6

15.7

15.8

HIMU

EMII

EMI

DMM

143 N

d/14

4 Nd

206Pb/204Pb

207 P

b/20

4 Pb

87Sr/86Sr

C

0.702 0.703 0.704 0.705 0.706

0.5124

0.5126

0.5128

0.5130

0.5132

0.5134 DMM

HIMU

EMI

EMII

D

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

B


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A

East Pacific Rise

Figure 11. Trace-element and isotopic data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) for basal-tic rocks from the East Pacifi c Rise mid-ocean ridge. DMM—depleted MORB mantle, HIMU—high μ, EMI—enriched mantle I, EMII—enriched mantle II.


cordance along the active spreading ridge and recognized several components in the mantle sources. Subsequent work (Pyle et al., 1995) used older ocean fl oor basalts to confi rm the existence of the Australian-Antarctica discordance during the opening of the ocean basin between Australia and Antarctica and to map its late Mesozoic-Cenozoic migration. The isotopic composition of Pacifi c-type MORB refl ects a mixture of DMM and HIMU mantle compo-nents (Pyle et al., 1992). The isotopic composition of Indian-type MORB is more complex and requires mixing of DMM, HIMU, and EMI components (Pyle et al., 1992). Trace-element data (Pyle et al., 1992, 1995) for the Australian-Antarctica discordance are shown in Figures 13A and 13B, both for active ridge (fi lled crosses) and older ocean fl oor (open crosses) basalts. On the Th/Y-Nb/Y ratio plot, the data plot along the entire mantle depletion-enrichment array, from values more depleted than N-MORB to values more enriched than E-MORB. One sample of young basalt has an elevated Th/Yb ratio similar to a subduction component. Trace-element data generally exhibit N-MORB and E-MORB pat-terns on the spider diagrams, and a few samples show positive U (but not Pb) values. Isotopic data shown in Figures 13C and 13D, only for Holocene (active ridge) basalts, are consistent with mixing of DMM, HIMU, and EMI components (Pyle et al., 1992).

Red Sea

The Red Sea provides an example of seafl oor spreading in a nascent ocean basin (younger than 5 Ma) formed by continen-tal rifting (Volker et al., 1997). Trace-element and isotopic data for basalts collected from the Red Sea region are shown in Fig-

ures 14A and 14B and include samples from the axial rift zone (Ramad seamount, Hanish-Zukir Islands) and from the fl anks of the ocean basin (Hamdan and Jizan volcanic fi elds). Nd-Sr-Pb isotope data provide evidence for a heterogeneous mantle source that includes mixing among several isotopic components (Figs. 14C–14D). Volker et al. (1997) argued for mixing among a MORB-type mantle (DMM), an Afar plume component (HIMU), and an EMI-EMII hybrid component derived from continental lithosphere. Trace-element data for the Red Sea samples show mantle enrichments similar to E-MORB and OIB on both the spider diagram and the Th/Y-Nb/Y ratio plot (Figs. 14A–14B).

ADDITIONAL EVIDENCE FOR A SUPRASUBDUCTION-ZONE ORIGIN FOR OPHIOLITES

Workers who prefer a mid-ocean-ridge origin for most or all ophiolites attempt to counter the suprasubduction-zone inter-pretation of ophiolite genesis by pointing to “...the inadequacy of geochemistry itself to determine the tectonic environment...” of individual an ophiolite (Moores, 2003, p. 26). Such criticism attempts to cast reasonable doubt on the validity of geochemical interpretations. However, the suprasubduction-zone interpretation is not based solely on geochemical data on volcanic rocks but rather on a range of petrologic and geologic data, of which geochemis-try is a major component. Next, we review three additional lines of evidence that support a suprasubduction-zone origin for much of the ophiolite record: evidence of wet magmas, the sedimentary cover on ophiolites, and the issue of ophiolite preservation.

South Atlantic Ridge

17.0 18.0 19.0 20.0 21.0

15.4

15.5

15.6

15.7

15.8

HIMU

EMII

EMI

DMM

207 P

b/20

4 Pb

206Pb/204Pb

143 N

d/14

4 Nd

87Sr/86Sr

C

0.702 0.703 0.704 0.705 0.706

0.5124

0.5126

0.5128

0.5130

0.5132

0.5134 DMM

HIMU

EMI

EMII

D


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

B

Figure 12. Trace-element and isotopic data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) for basal-tic rocks from the South Atlantic mid-ocean ridge. DMM—depleted MORB mantle, HIMU—high μ, EMI—enriched mantle I, EMII—enriched mantle II.


Evidence of Hydrous Magmas

Water plays a major role in petrologic models for the genesis and evolution of subduction-zone magmas, in contrast to water-poor environment at mid-ocean ridges. As discussed already, aqueous fl uids derived from the subducting slab carry LILEs into the mantle wedge source of suprasubduction-zone basalts, lower solidus temperatures in the source, and trigger melting. Fraction-ation of these wet basalts in the lithosphere produces crystalliza-tion sequences in arc basalts that are in contrast to those observed in MORBs. In addition, the concentration of water vapor in the residual magma during fractionation may lead to the separation of a hydrous vapor phase (retrograde boiling) and the formation of miarolitic cavities in isotropic gabbros of the upper plutonic series (Fig. 15A).

In MORB, the typical observed crystallization sequence is olivine/spinel > plagioclase > clinopyroxene ± ortho-pyroxene (Bryan, 1983; Pearce et al., 1984). In arc basalts, the typical observed crystallization sequence is olivine/spinel > clinopyroxene/orthopyroxene > plagioclase > hornblende (Pearce et al., 1984; Cameron, 1985). Experimental results for low-pressure (~2 kbar) crystallization of wet tholeiite (Sisson and Grove, 1993) confi rm the role of water in the production of the typical crystallization sequence in arc basalts. These experiments show that hydrous basalts crystallize olivine, clinopyroxene, and plagioclase (±spinel or magnetite); as magma chemistry evolves toward more siliceous composition, olivine, clinopyroxene, and An-rich plagioclase become unstable and react with the melt to form orthopyroxene, hornblende, and more Ab-rich plagio-

clase. Reaction textures with resorbed olivine, clinopyroxene, and An-rich plagioclase enclosed in hornblende oikocrysts are evidence of the crystallization of wet basalt and are reported in subduction-related magma systems.

In the ophiolite record, crystallization sequences can be determined from phenocryst assemblages in the volcanic section and from cumulate sequences in the plutonic sec-tion. In suprasubduction-zone ophiolites, basal cumulate sequences are dominated by dunite, wehrlite, and clino-pyroxenite overlain by pyroxene gabbro and hornblende gabbro (Fig. 15B). Hawkins and Evans (1983) reported well-displayed suprasubduction-zone cumulate sequences from the Zambales Range in the Luzon ophiolite. Thus, ophiolites with suprasubduction-zone geochemistry exhibit crystalliza-tion sequences (pyroxene before plagioclase) indicative of wet magmas (Table 3). Hornblende gabbro is a volumetri-cally important constituent of suprasubduction-zone ophio-lites and requires water-bearing magmas; in some cases, large, decimeter-scale hornblende crystallizes in the isotropic gabbros and diorites where water concentration is highest (Figs. 15C and 15D). This evidence for hydrous magmas in suprasubduction-zone ophiolites places the geochemical data in a broader petrologic context. This evidence is in sharp con-trast to crystallization sequences observed in modern MORB, where plagioclase appears before pyroxene and basal cumu-lates should be composed of dunite, troctolite, and anorthosite with hornblende as only a minor constituent (Pearce et al., 1984). The MORB ophiolites reviewed here exhibit dry crys-tallization (plagioclase before pyroxene) sequences (Table 3).

Australia - Antarctica Discordance

17.0 18.0 19.0 20.0 21.0

15.4

15.5

15.6

15.7

15.8

HIMU

EMII

EMI

DMM

207 P

b/20

4 Pb

143 N

d/14

4 Nd

206Pb/204Pb 87Sr/86Sr

C

0.702 0.703 0.704 0.705 0.706

0.5124

0.5126

0.5128

0.5130

0.5132

0.5134 DMM

HIMU

EMI

EMII

D


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

B

Figure 13. Trace-element and isotopic data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) for basal-tic rocks from the Australian-Antarctica discordance section of the Southeast In-dian mid-ocean ridge. DMM—depleted MORB mantle, HIMU—high μ, EMI—enriched mantle I, EMII—enriched mantle II.

Red Sea

17.0 18.0 19.0 20.0 21.0

15.4

15.5

15.6

15.7

15.8

HIMU

EMII

EMI

DMM

207 P

b/20

4 Pb

143 N

d/14

4 Nd

206Pb/204Pb 87Sr/86Sr

C

0.702 0.703 0.704 0.705 0.706

0.5124

0.5126

0.5128

0.5130

0.5132

0.5134 DMM

HIMU

EMI

EMII

D


Rb Th Nb La Y Lu

Rock/NMORB

100

10

1

0.1

0.01

A

NMORB

EMORB

OIB

0.01 0.1 1 10 100Nb/Yb

10

1

0.1

0.01

Th/Yb

B

BA

DC

Figure 14. Trace-element and isotopic data for normal (N) mid-ocean-ridge basalt (MORB), enriched (E) MORB, and ocean-island basalt (OIB) for basaltic rocks from the Red Sea mid-ocean ridge. DMM—depleted MORB mantle, HIMU—high μ, EMI—enriched mantle I, EMII—enriched mantle II.

Figure 15. Field photographs of plutonic rocks from ophiolites documenting “wet” magmas: (A) miarolitic cavities, Point Sal ophiolite, California, (B) cumulate clinopyroxenite and wehrlite, Point Sal ophiolite, California, (C) decimeter-scale hornblendes in appinite dike, Trinity ophiolite, California, (D) photomicrograph of quartz diorite, showing abundant quartz (clear), feldspar (pale brown, low relief), and hornblende (dark brown, high relief), Elder Creek ophiolite, Califor-nia, fi eld of view = 5.2 mm, plane light. Scale in A is 15 cm long; hammers in B and C are ~35 cm long.


Peridotite Mineral Chemistry: Evidence for Hydrous Melting

Peridotites form a signifi cant fraction of most ophiolite assemblages, but they are seldom as well studied as the crustal sections. There now exists a large body of data on both abyssal peridotites, dredged from oceanic fracture zones and other base-ment exposures, and suprasubduction-zone (forearc) peridotites, sampled from fault scarps on the leading edge of subduction com-plexes. Abyssal peridotites are dominantly lherzolite, consisting of olivine, enstatite, Cr-diopside, and aluminous spinel (e.g., Dick and Bullen, 1984; Dick, 1989; Johnson et al., 1990). In con-trast, forearc peridotites consist largely of harzburgite, consisting of olivine, enstatite, and Cr-rich spinel, and dunite, consist ing of olivine plus chromite (Ishii et al., 1992; Parkinson and Pearce, 1998; Pearce, 2003). In general, pyroxenes from abyssal perido-tites are relatively rich in incompatible elements compared to pyroxenes from forearc peridotites, which have extremely low incompatible element concentrations.

Perhaps the most useful mineral, however, is spinel, which varies systematically in composition in response to melt extraction and is resistant to low-temperature alteration during serpentini zation; it is commonly the only primary phase remain-ing in highly serpentinized peridotites. Abyssal perid otites are characterized by relatively aluminous spinels, with Cr# (100 × Cr/[Cr + Al]) ranging from ∼10 to 59 (Fig. 16), indi-cating limited melt extraction (Dick and Bullen, 1984; Dick, 1989). In contrast, peridotites dredged from forearc regions are characterized by relatively Cr-rich spinels, with Cr# rang-ing from ~38 to 60 in harzburgites and ∼60 to 84 in dunites (Fig. 16), indicating more extensive melt extraction in response to hydrous melting (Ishii et al., 1992; Parkinson and Pearce,

1998; Pearce, 2003). The extremely high Cr# observed in many forearc peridotite spinels is also characteristic of spinels in high-Mg andesites and dacites, and other boninitic lavas. In general, spinels from perido tites associated with back-arc basins (e.g., Mariana Trough) have compositions similar to those from abys-sal peridotites (e.g., Ohara et al., 2002).

Spinel compositional data are not widely available for many ophiolite mantle sections, but the data available suggest that ophiolite peridotites fall into two groups: those with clear suprasubduction-zone affi nities and those with mixed MORB–suprasubduction-zone affi nities. Vourinos is dominated by suprasubduction-zone spinel compositions (Cr# 45–85; Kon-stantopoulou and Economou-Eliopoulos, 1990). Troodos is similarly dominated by suprasubduction-zone spinel compo-sitions (Cr# 48–82; Hebert and Laurent, 1990; Georgiou and Xenophontos, 1990), but it also contains small domains of lherzo lite with spinel Cr# of 22–28 (Batanova and Sobolev, 2000). The Lewis Hills massif in the Bay of Islands complex resembles Troodos and is dominated by suprasubduction-zone spinels with Cr# of 50–78, but it contains a geographically small domain of abyssal peridotite with spinel Cr# of 15–30 (Suhr and Edwards, 2000). Oman shows the greatest variation, with peridotite spinel Cr# ranging from 21 to 67, where most values are greater than 40 (Le Mee et al., 2004).

The signifi cance of the mixed domains is unclear because our database for forearc peridotites is relatively small, but the conclu-sions reached from peridotite spinel compositions are generally consistent with those derived from volcanic rock geochemistry. Ophiolites with volcanic rocks that indicate a suprasubduction-zone origin are dominated by Cr-rich spinels with Cr# mostly >50 and few if any abyssal peridotite composition spinels. Ophiolites with mixed MORB–suprasubduction-zone volcanic sections are

0 10 20 30 40 50 60 70 80 90 100Cr#

0

20

40

60

80

100

Abyssal PeridotitesForearc LherzoliteForearc HarzburgiteForearc DuniteBoninite

Mg#

Abyssal Peridotites

Forearc Peridotites

BoninitesFigure 16. Spinel composition plot for abyssal peridotites, forearc peridotites, and boninites. Abyssal peridotites have low Cr# (100 × Cr/[Cr + Al]) compared to forearc peridotites and boninites. Data are from: abyssal peridotites—Dick and Bullen (1984), Dick (1989), Juteau et al. (1990), Komor et al. (1990), Hellebrand et al. (2002), Ohara et al. (2002), Arai and Matsukage (1998); forearc perid-otites—Parkinson and Pearce (1998), Ishii et al. (1992), Arai et al. (1990), Franz et al. (2002), Okamura et al. (2006); boninites—Falloon et al. (1989), Van Der Laan et al. (1992).


underlain by peridotites with a range of spinel compositions that refl ect both abyssal peridotite and forearc peridotite composi-tions, but even in these examples, high-Cr spinels characteristic of forearc peridotites are most common.

Sedimentary Cover

One major focus of the “ophiolite conundrum” is the cover of pelagic sediment, which is commonly, but not universally, associated with ophiolites. The crux of this argument is that the pelagic cover found on many ophiolites is inconsistent with for-mation of those ophiolites in or near an island-arc setting. The most common pelagic sediment associated with Mesozoic ophio-lites is chert; indeed, chert was part of “Steinmann’s Trinity” and until the 1972 Penrose conference on ophiolites was considered an essential part of ophiolite stratigraphy. These cherts are typi-cally rich in radiolaria and are thought to represent the slow accu-mulation of radiolarian ooze on the seafl oor over several million years prior to obduction (e.g., Pessagno et al., 2000).

Pure radiolarian cherts, with little or no clastic component, are generally restricted to the highly dismembered, incomplete fragments of oceanic crust found in accretionary complexes or some collision zones, for example, the Franciscan assemblage of California (Karl, 1984; Murchey, 1984; Murchey and Jones, 1984) and ophiolites of the Western Mediterranean (Alps, Apen-nines; Bill et al., 2001). These cherts contain abundant radio-laria and consist of nearly pure silica. They may also represent large time spans, e.g., Franciscan cherts of the Marin Headlands terrane in California, which represent ~30 m.y. of accumulation on the seafl oor (Murchey and Jones, 1984). Ophiolites in the northern Apennines apparently formed close to a rifted conti-nental margin (Rampone and Piccardo, 2000), so only a limited age range is expected.

In contrast, cherts overlying many suprasubduction-zone ophiolites are rich in volcanic ash, and many are essentially altered tuffs. Cherts associated with the Coast Range ophiolite in California contain up to 18% alumina and minor radiolaria (Hopson et al., 1981). In some Coast Range ophiolite locations, these altered tuffs are overlain by volcaniclastic sections up to 1.5 km thick with intercalated “radiolarian tuffs.” Similarly, radiolarites overlying the Troodos ophiolite contain common ash layers and grade upward into calc-alkaline volcaniclastic strata of the Kanaviou Formation (e.g., MacLeod et al., 1990). In Oman, tuffaceous chert overlies arc volcanics of the Alley unit and is overlain by ocean-island basalts of the Sahali volcanics. In the Josephine ophiolite, calc-alkaline volcaniclastic detritus is found as interpillow sediment in the volcanic section, showing its clear relationship to arc volcanism (Pessagno et al., 2000). In all of these cases, sediments deposited on the ophiolite contain signifi -cant arc-derived detritus.

Recent work in the western Pacifi c (Fryer et al., 2000; Hawkins, 2003) has shown that the extensional forearc envi-ronments thought to characterize suprasubduction-zone ophio-lite formation are generally not the locus of thick arc-derived

sedimentation. There are two reasons for this. First, extensional forearcs associated with nascent subduction zones tend to be wide (150–300 km) with diffuse extension distributed across the width. The crust is thin and broken into a series of linear horsts and grabens parallel to the trench axis, which will trap clastic sediment near its source and prevent its distribution throughout the forearc (Fryer et al., 2000; Hawkins, 2003). Second, nascent arcs lack a distinct volcanic front and an emergent arc edifi ce. As long as most volcanic activity occurs underwater, the distribu-tion of ash and coarse volcaniclastic materials will be limited.

Preservation and Emplacement

The structural preservation of ophiolites relates in large part to their emplacement mechanics. Moores (1998) grouped ophio-lites into two broad categories: Tethyan ophiolites and Cor dilleran ophiolites. Tethyan-type ophiolites (e.g., Oman, Troodos , Pindos, Vourinos, Muslim Bagh) are emplaced onto passive continental margins and are typically overlain by sediments characteristic of passive-margin settings (limestone, dolomite). In contrast, Cordilleran-type ophiolites (e.g., Coast Range ophiolite of Cali-fornia, Trinity ophiolite, Cape Vogel) are associated with active continental margins and typically are underlain by accretion-ary complexes and overlain by clastic sediments deposited in a forearc basin setting (turbidites, mudstones, conglomerates).

Tethyan- and Cordilleran-type ophiolites are similar in that both types commonly display complete or near complete ophio-lite stratigraphy (as defi ned by Penrose Conference Participants, 1972), are relatively intact structurally, and are characterized by suprasubduction-zone lava compositions (e.g., Shervais, 2001). Thus, we infer that both Tethyan- and Cordilleran-type ophiolites form in the upper plates of subduction zones and differ largely in their mode of emplacement (Fig. 17). Tethyan-type ophiolites represent obduction of forearc lithosphere onto a passive con-tinental margin during the attempted subduction of the passive margin (Figs. 17A–17D). A variation on the normal Tethyan type of ophiolite is found in the Alps, where the upper-plate ophiolite represents tectonically thinned continental lithosphere that has not been signifi cantly modifi ed by arc volcanism (Frisch et al., 1994). Cordilleran-type ophiolites are emplaced by “accretionary uplift” (Shervais, 2001), where continued growth of the under-lying accretionary complex gradually lifts the overlying forearc ophiolite assemblage—there is typically no collision with a pas-sive continental margin (Figs. 17E–17H).

We commonly fi nd ancient rock assemblages with geochemi-cal and petrologic characteristics that resemble true oceanic crust formed at mid-ocean-ridge spreading centers or intraplate oceanic islands as dismembered, incomplete fragments within subduction-zone accretionary complexes. Complete ophiolite sections are unknown in these complexes, and gabbro is rare, but volcanic rocks overlain by chert and associated with mantle-derived perido-tite (now serpentinite) are common. In all cases, the volcanic rocks associated with these complexes are geochemically equivalent to N-MORB, E-MORB, or OIB; arc-like suprasubduction-zone vol-


canics are rare or nonexistent. Examples of this category include fragments of oceanic crust found in the Franciscan assemblage (e.g., Shervais, 1990, 2006; MacPherson et al., 1990) and in the Apennines (Rampone and Piccardo, 2000).

These dismembered fragments represent ocean crust formed either at mid-ocean-ridge spreading centers or on off-axis seamounts or plateaus. Unlike suprasubduction-zone ophiolites, these ophiolite remnants are never complete stratigraphically. The scarcity of gabbro in accretionary com-plexes suggests that they may preferentially sample oceanic crust formed near fracture zones, where mantle serpentine is exposed on the seafl oor and volcanic rocks may be erupted directly onto serpentine, with no other intervening crust (e.g., Coleman, 2000). Off-axis seamounts or oceanic plateaus are not well anchored structurally to the underlying seafl oor, and they may also be preferentially detached during subduction

and emplaced within the accretionary complex. One characteris-tic of these Franciscan-style accretionary complexes is that they sample material with a wide range in ages and are assembled over a prolonged time period during continuous subduction of the subjacent oceanic lithosphere (Shervais, 2006).

A variation on the classic Franciscan-style accretionary com-plex is found in Cyprus and Oman, where alkali basalt seamounts and fringing reefs associated with rifting of the passive margin are scraped off the subducting plate during collision between the ophio lite and the passive margin. These detached seamounts are mixed into the adjacent passive-margin sediments to form a schüp-penzone beneath the ophiolite; examples include the Mamonia complex and Ayia Varvara Formation in Cyprus ( Malpas et al., 1992, 1993; Robertson and Xenophontos, 1993) and the Oman exotics within the Hawasina nappes (Robertson, 1986; Bechennec et al., 1988). These complexes sample seamounts that formed

A SSZ ophiolite forms over sinking slab

B Ophiolite collides with ridge crest, ophiolite formation stops

C Ophiolite encounters passive margin, begins to thrust over sediment wedge

D Ophiolite is emplaced onto passive margin above schüppenzone of passive-margin sediments

E SSZ ophiolite forms over sinking slab.

F Ophiolite collides with ridge crest, ophiolite formation stops.

G Continued subduction; formation of accretionary complex uplifts ophiolite.

H Growth of accretionary complex exhumes ophiolite from beneath cover of forearc sediments.

Figure 17. Cross-section models of ophiolite emplacement by (A–D) obduction and (E–H) accretionary uplift. (A) Suprasubduction-zone (SSZ) ophiolite forms over sinking slab, which is separated from a passive continental margin by a spreading center. (B) Ophiolite encounters the spreading center; ophiolite formation stops, and the basal part of the ophiolite is thermally metamorphosed by high heat fl ux from the thin litho-sphere near the ridge crest. (C) Ophiolite encounters sedimentary wedge of the passive margin, which is depressed below the ophiolite and over-ridden by it; imbricate thrust sheets form in the passive-margin sediments. (D) Ophiolite is emplaced onto the passive continental margin above a schüppenzone of imbricate thrust sheets in the passive-margin sediments (e.g., Hawasina nappes in Oman, Mamonia complex in Cyprus). (E) Suprasubduction zone (SSZ) ophiolite forms over sinking slab; sinking of slab slows as spreading center is approached. (F) Ophiolite en-counters the spreading center; ophiolite formation stops, and the basal part of the ophiolite is thermally metamorphosed by high heat fl ux from the thin lithosphere near the ridge crest. (G) Sediments deposited in the subduction-zone trench are subducted to form an accretionary prism beneath the leading edge of the ophiolite; abyssal sediments and volcanic rocks scraped off the subducting oceanic plate may be included in the accretionary prism, which is dominated by juvenile detritus from the upper plate. (H) The accretionary prism continues to grow and thicken, exhuming the leading edge of the ophiolite.


over a limited time period during initial rifting of the passive mar-gin; the complexes themselves were formed during closure of the ocean basin that was created during this rifting event.

Sturm et al. (2000) proposed that oblique subduction of a spreading center will result in suprasubduction zone–type enrich-ments of the spreading center magmas due to sublithospheric fl ow through the “slab window,” as discussed earlier (ridge-trench-trench junction model). They also suggested that oblique subduction of the two plates bordering the spreading center cre-ates a setting favorable for ophiolite emplacement, as shown by the Taitao ophiolite complex (Stern, 1980). In this model, the fi rst slab to be subducted (on the trench side of the spreading center) is stranded in the subduction zone, and the second slab (from the opposite side of the spreading center) is subducted beneath it—effectively stepping the subduction zone farther outboard from the arc. They further suggested that this model may apply to suprasubduction-zone ophiolites like Semail and Bay of Islands (Sturm et al., 2000).

A geochemical assessment of this model has already been presented; we present here a geometric assessment. This model implies that the ophiolite will be emplaced structurally beneath a pre-existing island arc and may be separated from the overlying arc by an accretionary complex. Should this arc collide with and subduct a passive margin (e.g., Tethyan ophiolites), the ophio-lite will be a small part of the total package; the older arc and its accretionary complex will dominate. This is not observed in Semail, Troodos, or any other Tethyan-type ophiolites; in contrast, these ophiolites are overlain stratigraphically by post-collisional platform sediments (e.g., Glennie et al., 1974). There is no evidence for an older arc complex that structurally overlies the “pseudo-suprasubduction-zone” ophiolite.

The ridge-trench-trench model may apply to some Cordilleran-type ophiolites, which often have older arc complexes behind them. Unfortunately, most Cordilleran-type ophio lites are overlain by thick accumulations of forearc basin sediment, which obscures primary tectonic relationships. In this case, detailed chemical/petrologic studies must be applied.

HISTORICAL CONTINGENCY REDUX

At the beginning of this paper, we presented a summary of the main precepts of the historical contingency model, as pro-posed by Moores et al. (2000). In this section, we assess the applicability of each of these precepts to ophiolite generation, guided by our exploration of modern tectonic settings and the rocks that form them.

Asthenosphere Modifi ed by a Previous Subduction Event

This represents one of the central precepts of the historical contingency model. The only clear example we have of this process in the recent geologic past is the Woodlark basin in the southwest Pacifi c. Collision of the Ontong Java plateau with the Solomon arc along a SW-dipping subduction zone stalled subduction of the

Pacifi c plate and forced the inception of a new NE-dipping subduc-tion zone that consumed the former back-arc basin. As discussed previously, the geochemistry of basalts from the Woodlark basin is typical of back-arc basin basalts: they are generally MORB-like in composition but have a faint subduction-zone signature in the more mobile large ion lithophile elements, such as Pb.

In any case, basalts formed in such a setting still face the problem of emplacement: how do you move dense rocks from the lower plate of a subduction zone onto a passive margin (obduction) or place them above an accretionary prism (accre-tionary uplift) without violating the laws of physics? Like true mid-ocean-ridge basalts, back-arc basin basalts that are sub-ducted beneath their parent arc due to a subduction polarity fl ip are unlikely to be preserved, except as small fragments and slivers within the accretionary complex of the subduction zone.

Formation of a Slab Window Where Ridges Are Subducted Orthogonally or Obliquely

The Chile and Juan de Fuca spreading ridges are modern examples of this process. As discussed earlier, in both of these ridges, the dominant geochemical signature is that of E-MORB enrichment, which is unrelated to their position adjacent to a subduction zone. Subduction enrichment in large ion litho-phile elements is minimal and superimposed on the dominant E-MORB enrichment. Thus, while this process may inject small volumes of subduction-enriched mantle into the spreading cen-ter, it is not suffi cient to create the dominant pattern documented here of depletion in the more incompatible elements relative to N-MORB, strong negative anomalies in Nb and the other high fi eld strength elements, and signifi cant enrichments in the fl uid-mobilized large ion lithophiles/low fi eld strength elements.

Dynamically, there is some merit to the suggestion that sub-duction of a spreading axis may allow ridge segments to be more easily emplaced onto a continental margin: the spreading axis forms a discontinuity that allows the more distal plate to be thrust under the more proximal plate (which enters the subduction zone fi rst), potentially trapping the proximal plate above a newly con-fi gured subduction boundary. Geometrically, this model implies that the trapped portion of the proximal plate will be preserved beneath a volcanic arc and its previous accretionary complex. Few, if any, ophiolites preserve this geometry: for example, the Coast Range ophiolite of California lies above the Franciscan accretionary complex, not below it, and it is overlain deposition-ally by arc-derived sediments. Seismic tomography shows that during most ridge-trench collisions, both sides of the spreading axis are subducted and sink into the mantle (e.g., Rogers et al., 2002). If fragments of oceanic crust are emplaced in this way, it is probable that they will be preserved within the accretionary complex as large mélange blocks. It is more likely that the slab window allows MORB-source asthenosphere to affect the mantle wedge above the subduction zone, infl uencing the compositions of the resulting arc magmas (Shervais et al., 2004, 2005a, 2005b; Sisson et al., 2003, and papers therein).


Subducted Slab Component—The Fate of Old Plates

Numerical models of mantle dynamics suggest that chemical and isotopic heterogeneities related to subduction of oceanic litho-sphere may persist on extremely long time scales (e.g., Kellogg et al., 1999). It has been known for some time that subducted slabs may remain distinct parts of the mantle, with chemical and isotopic systematics that differ from the surrounding MORB-source astheno-sphere. These slabs have been imaged by seismic tomography as cooler, higher-velocity regions in the mantle (Rogers et al., 2002). However, these subducted slabs do not represent the subduction component that modifi es the source region of suprasubduction-zone magma systems (arcs and ophiolites). Lithospheric slabs that are subducted deep into the mantle represent the residues that remain after extraction of the fl uid-rich component that carries silica, alka-lis, and the large ion lithophile/low fi eld strength elements into the mantle wedge above the subduction zone. This residual slab com-ponent is enriched relatively in the incompatible elements that are not mobilized in high-temperature fl uids, i.e., the high fi eld strength elements such as Ti, Nb, Ta, Hf, and Zr.

Trace-element systematics suggest that many OIBs form largely by the remelting of recycled oceanic lithosphere and depleted MORB-source asthenosphere (e.g., Hofmann, 1982; Weaver, 1991). In addition, ridge-centered oceanic islands like Iceland and the Azores infl uence the composition of spreading center basalts by introducing these same components into the melting zone by fl ow along sublithospheric conduits (Schilling, 1973). As a result, basalts erupted along mid-oceanic ridges vary in chemical and isotopic composition from “normal” N-MORB to “enriched” E-MORB to true OIB at the ridge-centered oceanic islands. In this context, the OIB-style “enriched” basalt refers to a general enrichment in incompatible trace elements including both LILE and HFSE, not to the fl uid-mobilized, LILE enrich-ment seen in suprasubduction-zone processes. This distinction is important, because much of the historical contingency model rests on the defi nition of enrichment. In subduction-zone enrichment, there is a strong decoupling between LILEs (which are mobilized by aqueous solution and transferred from the subducted slab to the mantle wedge) and HFSEs (which are insoluble in hydrous fl uids and remain in the slab). The OIB-style enrichment is mobi-lized by silicate melts; LILE and HFSE are not decoupled, but the enrichment sources (old slabs) were previously depleted in LILE (subduction) components. In effect, the enriched component found in OIB and E-MORB is the result of subduction, but it rep-resents the complement to that found in the fl uid-fl uxed mantle wedge above a subduction zone. The processes that affect the mantle wedge defi ne what most geologists refer to as subduction-zone enrichment.

Isotopic Components in Ocean-Island Basalts

Ocean-island basalts (OIBs) contain a number of distinct isotopic components (DMM, HIMU, EM1, EM2, PREMA/FOZO) that represent the persistence of long-term trace-element

heterogeneities in the mantle (Zindler and Hart, 1986; Hart, 1988) generated by subduction recycling. As discussed already, the long-term trace-element heterogeneities that are responsible for the isotopic components found in OIB and MORB result from the decoupling of incompatible trace elements that are mobilized in hydrous slab-derived fl uids to fertilize the overly-ing mantle wedge from those elements that are not mobilized in high-temperature hydrous fl uids and thus remain behind in the slab. Some of these components may represent sediments that are carried deep into the subduction zone and recycled into the mantle, but all differ from the short-term enrichments in silica, alkalis, and low fi eld strength elements that are characteristic of suprasubduction magma systems. Thus, the fact that MORB and OIB preserve isotopic evidence for a range of trace-element enrichment processes is irrelevant to any discussion of the origin of suprasubduction-zone magmas, except where these sources may be trapped in the mantle wedge above a subduction zone and participate in the formation of arc-related magmas.

Models for Mid-Ocean-Ridge Processes

The historical contingency model requires us to believe that the only oceanic crust preserved intact is that formed over previ-ously modifi ed lithosphere; ocean crust formed from normal or plume-enriched MORB asthenosphere that has not been modifi ed by these cryptic subduction-like enrichment processes is not pre-served as ophiolites, even though >98% of all oceanic crust today is normal or plume-enriched MORB with less than 52 wt% silica.

The historical contingency model uses isotopic hetero-geneities in the modern mantle refl ected in data from OIB, coupled with models of mantle dynamics, to argue that during certain periods of Earth history, mid-ocean-ridge magma systems could have tapped subduction-modifi ed mantle, yielding basalts with supra subduction zone–like compositions. Implicit in this argument is the assumption that the present is not one of those periods of Earth history; in other words, modern mid-ocean ridges presently do not tap the type of subduction-modifi ed mantle represented by OIB isotopic data. Our review of recently published trace-element and isotopic data from active mid-ocean ridges, however, refute this implicit assumption. Although the modern MORB data set refl ects the same isotopic variations observed in OIB, variations attributed to contributions from subducted oceanic lithosphere, these mid-ocean-ridge basalts do not exhibit the suprasubduction zone–like, LILE enrichments seen in either modern subduction-zone basalts or in the suprasubduction-zone ophiolite record. What these basalts do show are OIB-style enrichments, i.e., a coupled LILE-HFSE enrichment, as discussed already.

DISCUSSION

The issues raised by the hypothesis of “historical contin-gency” and our discussion here are not trivial. What is at stake? Our reconstructions of global tectonics before the current ocean basins formed depend critically on how we interpret the ophio-


lite record. Even suprasubduction-zone ophiolites require the existence of an ocean basin for their formation, but the details of mountain-building events and relations between allochthonous sheets in complex orogens hinge on whether a given ophio lite assemblage represents true “oceanic crust” formed at a mid-ocean-ridge spreading center, back-arc basin crust formed behind an arc, or suprasubduction-zone crust formed over a nascent sub-duction zone. These issues are particularly important for people who are not involved in the debate on how ophiolites form, but who use ophiolites to reconstruct the tectonic history of oro-gens. How do they know which model of ophiolite formation is correct, and how will it affect their interpretation of ophiolite assemblages in their fi eld area? Researchers who are not involved in this debate may be easily misled by hypotheses that appear sound, but which do not have the data to sustain them. The ability to use suprasubduction-zone ophiolites as natural laboratories for studying the initial composition of new subduction zones is also at stake; much of that record is found in suprasubduction-zone ophiolites formed at nascent forearc settings (Stern, 2004).

We must ask whether the term “ophiolite” is descriptive or genetic. If it is descriptive, then rock assemblages that fi t the description may form from a variety of processes, and the evidence for these processes can be applied to any assemblage that fi ts the description. If it is genetic, then we must thoroughly understand the origin of any given assemblage before we can con-sider whether this term applies. If we take the genetic approach, then most of the rock suites we call “ophiolites” would have to be renamed, because their petrology, geochemistry, and struc-tural setting are at odds with the interpretation that they represent oceanic crust formed at a mid-ocean-ridge spreading center.

We return to our suggestion that a new formulation of the “ophiolite conundrum” is needed: Given that many, if not most, ophiolites have geochemical signatures consistent with formation above active subduction zones, under what circumstances does this setting result in rock associations and structures consistent with those observed in ophiolites? Recent work in volcanic arcs of the western Pacifi c has shown that intra-arc rifting of existing primitive arcs (e.g., Vanuatu) and rifting associated with the onset of subduction (e.g., the Eocene Mariana-Izu-Bonin system) may form rock assemblages similar to those observed in ophiolite complexes (Bloomer et al., 1995; Hawkins, 2003; Fryer et al., 2000). These assemblages may represent up to 100% extension of the pre-existing crust over broad areas and may be starved of clastic sediments other than windborne volcanic ash. They also consist of volcanic and plutonic assemblages that have the petro-logic and geochemical characteristics of suprasubduction-zone magmas, including boninites, which appear to form only within the forearc regions of highly extended arcs.

Finally, we emphasize that it is not just the geochemistry of ophiolite volcanic rocks that requires their formation in supra-subduction environments. The evidence for hydrous magmatic systems, hydrous melting, and the common occurrence of vol-canic ash in overlying cherts are consistent with an original suprasubduction-zone setting for these ophiolites and are incon-

sistent with formation at a mid-oceanic spreading center. More importantly, the structural setting of ophiolites during emplace-ment requires that they formed in the upper plate of a convergent-margin system. In fact, based on their structural setting alone, we would be forced reach this conclusion (e.g., Gealey, 1977).

ACKNOWLEDGMENTS

The authors have benefi ted over the years from spirited dis-cussions with many ophioliteologists, but none has challenged us more than Cliff Hopson and Eldridge Moores, who forced us to examine our assumptions and think clearly about what is observed and what we infer. We are grateful to Jim Hawkins and Bob Stern, who provided thoughtful reviews that aided us in improving the manuscript, and editorial handling by Jim Wright.

REFERENCES CITED

Alabaster, T., Pearce, J.A., and Malpas, J., 1982, The volcanic stratigra-phy and petrogenesis of the Oman ophiolite complex: Contributions to Mineralogy and Petrology, v. 81, no. 3, p. 168–183, doi: 10.1007/BF00371294.

Arai, S., and Matsukage, K., 1998, Petrology of a chromitite micropod from the Hess Deep, equatorial Pacifi c: A comparison between abyssal and alpine-type podiform chromitites: Lithos, v. 43, p. 1–14, doi: 10.1016/S0024-4937(98)00003-6.

Arai, S., Ito, M., Nakayama, N., and Masuda, F., 1990, A suspect serpentinite mass in the Tokyo Bay area: Petrology and provenance of serpentinite pebbles in the Upper Cenozoic system in the Boso Peninsula, central Japan: Journal of Geological Society of Japan, v. 96, p. 171–179.

Arculus, R.J., and Powell, R., 1986, Source component mixing in the regions of arc magma generation: Journal of Geophysical Research, v. 91, p. 5913–5926.

Bach, W., Erzinger, J., Dosso, L., Bollinger, C., Bougault, H., Etoubleau, J., and Sauerwein, J., 1996, Unusually large Nb-Ta depletions in the North Chile ridge basalts at 36°50′ to 38°S: Major element, trace element and isotopic data: Earth and Planetary Science Letters, v. 142, p. 223–240, doi: 10.1016/0012-821X(96)00095-7.

Baragar, W.R.A., Varga, R.J., Verosub, K.L., Baglow, N., and Gibson, I.L., 1990, The sheeted dyke zone in the Troodos ophiolite, in Malpas, J., Moores, E.M., Panayiotou, A., and Xenophontos, C., eds., Ophiolites: Oceanic Crustal Analogues: Proceedings of the Symposium “Troodos 1987”: Cyprus, Nicosia, Geological Survey Department, p. 37–52.

Barth, M.G., Mason, P.R.D., Davies, G.D., Dijkstra, A.H., and Drury, M.R., 2003, Geochemistry of the Othris ophiolite, Greece: Evidence for refertilization?: Journal of Petrology, v. 44, p. 1759–1785.

Batanova, V.G., and Sobolev, A.V., 2000, Compositional heterogeneity in subduction-related mantle peridotites, Troodos massif, Cyprus: Geology, v. 28, no. 1, p. 55–58, doi: 10.1130/0091-7613(2000)28<55:CHISMP>2.0.CO;2.

Beccaluva, L., Coltorti, M., Giunta, G., and Siena, F., 2004, Tethyan vs. Cor-dilleran ophiolites: A reappraisal of distinctive tectono-magmatic features of supra-subduction complexes in relation to subduction mode: Tectono-physics, v. 393, p. 163–174.

Bechennec, F., Le Metour, J., Raby, D., Villey, M., and Beurrier, M., 1988, The Hawasina Basin: A fragment of a starved passive continental margin, thrust over the Arabian platform during obduction of the Semail Nappe: Tec-tonophysics, v. 151, p. 323–343, doi: 10.1016/0040-1951(88)90251-X.

Bedard, J.H., 1999, Petrogenesis of boninites from the Betts Cove ophio-lite, Newfoundland, Canada; identifi cation of subducted source com-ponents: Journal of Petrology, v. 40, p. 1853–1889, doi: 10.1093/petrology/40.12.1853.

Bedard, J.H., Lauziere, K., Tremblay, A., and Sangster, A., 1998, Evidence for forearc seafl oor-spreading from the Betts Cove ophiolite, Newfound-land: Oceanic crust of boninitic affi nity: Tectonophysics, v. 284, no. 3–4, p. 233–245, doi: 10.1016/S0040-1951(97)00182-0.


Bill, M., O’Dogherty, L., Guex, J., Baumgartner, P.O., and Masson, H., 2001, Radiolarite ages in Alpine-Mediterranean ophiolites: Constraints on the oceanic spreading and the Tethys-Atlantic connection: Geological Society of America Bulletin, v. 113, p. 129–143, doi: 10.1130/0016-7606(2001)113<0129:RAIAMO>2.0.CO;2.

Bloomer, S.H., and Hawkins, J.W., 1983, Gabbroic and ultramafi c rocks from the Mariana Trench, in Hayes, D.E., ed., The Tectonic Evolution of the Southwest Asian Seas and Islands, Part 2: American Geophysical Union Memoir 27, p. 294–317.

Bloomer, S.H., Taylor, B., MacLeod, C.J., Stern, R.J., Fryer, P., Hawkins, J.W., and Johnson, L., 1995, Early arc volcanism and the ophiolite problem: A perspective from drilling in the western Pacifi c, in Taylor, B., and Nat-land, J., eds., Active Margins and Marginal Basins of the Western Pacifi c: American Geophysical Union Geophysical Monograph 88, p. 1–30.

Bryan, W.G., 1983, Systematics of modal phenocryst assemblages in sub-marine basalts: Petrological implications: Contributions to Mineralogy and Petrology, v. 83, p. 62–74, doi: 10.1007/BF00373080.

Cameron, W.E., 1985, Petrology and origin of primitive lavas from the Troodos ophiolite, Cyprus: Contributions to Mineralogy and Petrology, v. 89, no. 2–3, p. 239–255, doi: 10.1007/BF00379457.

Cann, J.R., 2003, The Troodos ophiolite and the upper oceanic crust: A recipro-cal traffi c in scientifi c ideas, in Dilek, Y., and Newcomb, S., eds., Ophio-lite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 309–321.

Coleman, R.G., 2000, Prospecting for ophiolites along the California conti-nental margin, in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 351–364.

Cousens, B.L., Allan, J.F., Leybourne, M.I., Chase, R.L., and Van Wagoner, N., 1995, Mixing of magmas from enriched and depleted mantle sources in the northwest Pacifi c: West Valley segment, Juan de Fuca Ridge: Contri-butions to Mineralogy and Petrology, v. 120, p. 337–357.

Crawford, A.J., and Falloon, T.J., 1989, Experimental determination of boninite petrogenesis, in Crawford, A.J., ed., Boninites: London, UK, Unwin Hyman, p. 100–120.

Crook, K.A.W., and Taylor, B., 1994, Structure and Quaternary tectonic history of the Woodlark triple junction region, Solomon Islands: Marine Geo-physical Researches, v. 16, no. 1, p. 65–89, doi: 10.1007/BF01812446.

Davidson, J.P., 1996, Deciphering mantle and crustal signatures in subduction zones magmatism, in Bebout, G.E., School, D.W., Kirby, S.H., and Platt, J.P., eds., Subduction Top to Bottom: American Geophysical Union Geo-physical Monograph 96, p. 251–262.

Deschamps, A., and Lallemand, S., 2003, Geodynamic setting of Izu-Bonin-Mariana boninites, in Larter, R.D., and Leat, P.T., eds., Intra-Oceanic Sub-duction Systems: Tectonic and Magmatic Processes: Geological Society [London] Special Publication 219, p. 163–186.

Dick, H.J.B., 1989, Abyssal peridotites, very slow spreading ridges and ocean ridge magmatism, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the Ocean Basins: Geological Society [London] Special Publication 42, p. 71–105.

Dick, H.J.B., and Bullen, T., 1984, Chromian spinel as a petrogenetic indica-tor in abyssal and alpine-type peridotites and spatially associated lavas: Contributions to Mineralogy and Petrology, v. 86, p. 54–76, doi: 10.1007/BF00373711.

Dilek, Y., 2003, Ophiolite concept and its evolution, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geo-logical Society of America Special Paper 373, p. 1–16.

Einaudi, F., Pezard, P.A., Cocheme, J.J., Coulon, C., Laverne, C., and Godard, M., 2000, Petrography, geochemistry and physical properties of a continuous extrusive section from the Sarami Massif, Semail ophio-lite: Marine Geophysical Researches, v. 21, no. 3–4, p. 387–407, doi: 10.1023/A:1026752415989.

Elliott, T., Plank, T., Zindler, A., White, W., and Bourdon, B., 1997, Element transport from slab to volcanic front at Mariana arc: Journal of Geophysi-cal Research, v. 102, p. 14,991–15,019, doi: 10.1029/97JB00788.

Ernewein, M., Pfl umio, C., and Whitechurch, H., 1988, The death of an accre-tion zone as evidenced by the magmatic history of the Sumail ophiolite (Oman), in Boudier, F., and Nicolas, A., eds., The Ophiolites of Oman: Tectonophysics, v. 151, p. 247–274.

Ewart, A., Bryan, W.B., Chappell, B.W., and Rudnick, R.L., 1994, Regional geochemistry of the Lau-Tonga arc and backarc systems, in Hawkins, J.,

Parsons, L., Allan, J., et al., Proceedings of the Ocean Drilling Program, Scientifi c Results, Volume 135: College Station, Texas, Ocean Drilling Program, p. 385–425.

Falloon, T.J., and Danyushevsky, L.V., 2000, Melting of refractory mantle at 1.5, 2 and 2.5 GPa under anhydrous and H2O-undersaturated conditions; implications for the petrogenesis of high-Ca boninites and the infl uence of subduction components on mantle melting: Journal of Petrology, v. 41, no. 2, p. 257–283, doi: 10.1093/petrology/41.2.257.

Falloon, T.J., Green, D.H., and McCulloch, M.T., 1989, Petrogenesis of high-Mg and associated lavas from the North Tonga Trench, in Crawford, A.J., ed., Boninites and Related Rocks: London, UK, Unwin Hyman, p. 357–395.

Flower, M.F.J., 2003, Ophiolites, historical contingency and the Wilson cycle, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 111–135.

Franz, L., Becker, K.-P., Kramer, W., and Herzig, P.M., 2002, Meta somatic mantle xenoliths from the Bismarck microplate (Papua New Guinea)—Thermal evolution, geochemistry and extent of slab-induced meta-somatism: Journal of Petrology, v. 43, p. 315–343, doi: 10.1093/petrology/43.2.315.

Frisch, W., Ring, U., Duerr, S., Borchert, S., and Biehler, D., 1994, The Arosa zone and Platta nappe ophiolites (eastern Swiss Alps); geochemical char-acteristics and their meaning for the evolution of the Penninic Ocean: Jahrbuch der Geologischen Bundesanstalt Wien, v. 137, no. 1, p. 19–33.

Fryer, P., Lockwood, J.P., Becker, N., Phipps, S., and Todd, C.S., 2000, Signifi -cance of serpentinite mud volcanism in convergent margins, in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 35–52.

Gass, I.G., 1968, Is the Troodos massif of Cyprus a fragment of Meosozoic ocean fl oor?: Nature, v. 220, p. 32–42.

Gass, I.G., Lippard, S.J., and Shelton, A.W., eds., 1984, Ophiolites and Oceanic Lithosphere: Geological Society [London] Special Publication 13, 413 p.

Gealey, W.K., 1977, Ophiolite obduction and geologic evolution of the Oman Mountains and adjacent areas: Geological Society of America Bulletin, v. 88, p. 1183–1191, doi: 10.1130/0016-7606(1977)88<1183:OOAGEO>2.0.CO;2.

Georgiou, E., and Xenophontos, C., 1990, Chromite occurrences and associ-ated plutonic rocks in the Akapnou Forest, in Malpas, J., Moores, E.M., Panayiotou, A., and Xenophontos, C., eds., Ophiolites: Oceanic Crustal Analogues: Proceedings of the Symposium “Troodos 1987”: Cyprus, Nicosia, Geological Survey Department, p. 585–592.

Glennie, K.W., Boeuf, M.G., Clarke, M.W.H., Moody-Stuart, M., Pilaar, W.F.H., and Reinhardt, B.M., 1974, Geology of the Oman Mountains: Verhan-delingen van het Koninklijk Nederlands Geologisch Mijnbouwkundig Genootschap, v. 31, 423 p.

Harper, G.D., 2003a, Fe-Ti basalts and propagating-rift tectonics in the Jose-phine ophiolite: Geological Society of America Bulletin, v. 115, no. 7, p. 771–787.

Harper, G.D., 2003b, Tectonic implications of boninite, arc tholeiite, and MORB magma types in the Josephine ophiolite, California, Oregon, in Dilek, Y., and Robinson Paul, T., eds., Ophiolites in Earth History: Geo-logical Society [London] Special Publication 218, p. 207–230.

Hart, S.R., 1988, Heterogeneous mantle domains: Signatures, genesis and mix-ing chronologies: Earth and Planetary Science Letters, v. 90, p. 273–296.

Hawkins, J.W., 2003, Geology of suprasubduction zones—Implications for the origin of ophiolites, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 227–268.

Hawkins, J.W., and Evans, C.A., 1983, Geology of the Zambales Range, Luzon, Philippines: Ophiolite derived from an island arc-backarc basin pair, in Hayes, D.E., ed., The Tectonic Evolution of the Southwest Asian Seas and Islands, Part 2: American Geophysical Union Memoir 27, p. 124–138.

Hawkins, J.W., Bloomer, S.H., Evans, C.A., and Melchior, J.T., 1984, Evolu-tion of intra-oceanic arc-trench systems, in Carlson, R.L., and Kobayashi, K., eds., Geodynamics of Back-Arc Regions: Tectonophysics, v. 102 p. 175–205.

Hebert, R., and Laurent, R., 1990, Mineral chemistry of the plutonic section of the Troodos ophiolite; new constraints for genesis of arc-related ophio-lites, in Malpas, J., Moores, E.M., Panayiotou, A., and Xenophontos, C.,


eds., Ophiolites: Oceanic Crustal Analogues: Proceedings of the Sympo-sium “Troodos 1987”: Cyprus, Nicosia, Geological Survey Department, p. 149–163.

Hellebrand, E., Snow, J.E., and Mühe, R.K., 2002, Mantle melting beneath Gakkel Ridge (Arctic Ocean), abyssal peridotite spinel com-positions: Chemical Geology, v. 182, p. 227–235, doi: 10.1016/S0009-2541(01)00291-1.

Hess, H.H., 1955, Serpentinites, orogeny and epeirogeny, in Poldervaart, A., ed., Crust of the Earth—A Symposium: Geological Society of America Special Paper 62, p. 391–408.

Hickey, R.L., and Frey, F.A., 1982, Geochemical characteristics of boninite series volcanics; implications for their source: Geochimica et Cosmochimica Acta, v. 46, no. 11, p. 2099–2116, doi: 10.1016/0016-7037(82)90188-0.

Hofmann, A.W., 1982, Mantle plumes from ancient oceanic crust: Earth and Planetary Science Letters, v. 57, p. 421–436, doi: 10.1016/0012-821X(82)90161-3.

Hofmann, A.W., 1997, Early evolution of the continents: Science, v. 275, no. 5229, p. 498–499.

Hopson, C.A., Mattinson, J.M., and Pessagno, E.A., Jr., 1981, Coast Range ophiolite, western California, in Ernst, W.G., ed., The Geotectonic Devel-opment of California, Rubey Volume I: Englewood Cliffs, New Jersey, Prentice-Hall, p. 418–510.

Ishii, T., Robinson, P.T., Mackawa, H., and Fiske, R.S., 1992, Petrological studies of peridotites from diapiric serpentine seamounts in the Izu-Ogasawara-Mariana forearc, Leg 125, in Fryer, P., Pearce, J.A., and Strokking, L.B., eds., Proceedings of the Ocean Drilling Program, Sci-entifi c Results, Volume 125: College Station, Texas, Ocean Drilling Pro-gram, p. 445–486.

Ishikawa, T., Nagaishi, K., and Umino, S., 2002, Boninitic volcanism in the Oman ophiolite: Implications for thermal condition during transition from spreading ridge to arc: Geology, v. 30, no. 10, p. 899–902, doi: 10.1130/0091-7613(2002)030<0899:BVITOO>2.0.CO;2.

Jacobsen, S.B., and Wasserburg, G.J., 1980, Sm-Nd isotopic evolution of chondrites: Earth and Planetary Science Letters, v. 50, p. 139–155, doi: 10.1016/0012-821X(80)90125-9.

Jenner, G.A., Dunning, G.R., Malpas, J., Brown, M., and Brace, T., 1991, Bay of Islands and Little Port complexes, revisited; age, geochemical and iso-topic evidence confi rm suprasubduction-zone origin: Canadian Journal of Earth Sciences, v. 28, no. 10, p. 1635–1652.

Johnson, K.M., Dick, H.J.B., and Shimizu, N., 1990, Melting in the oceanic upper mantle; an ion microprobe study of diopside in abyssal peridotite: Journal of Geophysical Research, v. 96, no. B3, p. 2661–2678.

Johnson, R.W., Jaques, A.L., Langmuir, C.H., Perfi t, M.R., Staudigel, H., Dunkley, P.N., Chappell, B.W., Taylor, S.R., and Baekisapa, M., 1987, Ridge subduction and forearc volcanism; petrology and geochemistry of rocks dredged from the western Solomon arc and Woodlark basin, in Taylor, B., and Exon, N.F., eds., Marine Geology, Geophysics and Geo-chemistry of the Woodlark Basin–Solomon Islands: Menlo Park, Cali-fornia, Circum-Pacifi c Council for Energy and Mineral Resources, Earth Science Series, p. 155–226.

Juteau, T., Berger, E.T., and Cannat, M., 1990, Serpentinized, residual mantle peridotites from the Mar Median Valley, ODP Hole 670a (21°10N, 45°02W, Leg 109): Primary Mineralogy and Geothermometry, in Pro-ceedings of the Ocean Drilling Program, Scientifi c Results, Volume 106/109: College Station, Texas, Ocean Drilling Program, p. 27–45.

Kamenetsky, V.S., Everard, J.L., Crawford, A.J., Varne, R., Eggins, S.M., and Lanyon, R., 2000, Enriched end-member of primitive MORB melts: Petrology and geochemistry of glasses from Macquarie Island (SW Pacifi c): Journal of Petrology, v. 41, p. 411–430, doi: 10.1093/petrology/41.3.411.

Karl, S.M., 1984, Sedimentologic, diagenetic, and geochemical analysis of Upper Mesozoic ribbon cherts from the Franciscan assemblage at the Marin Headlands, California, in Blake, C., Jr., ed., Franciscan Geology of Northern California: Los Angeles, California, Society of Economic Paleon tologists and Mineralogists, Pacifi c Section, Field Trip Guide-book 43, p. 71–88.

Karsten, J.L., Klein, E.M., and Sherman, S.B., 1996, Subduction zone geochemical characteristics in ocean ridge basalts from the southern Chile Ridge; implications of modern ridge subduction systems for the Archean, in Ludden, J.N., Arndt, N.T., and Francis, D., eds., Evolu-tion of Mafi c Magmatism through Time: St. Malo, France, Lithos, p. 143–161.

Kellogg, L.H., Hager, B.H., and van der Hilst, R.D., 1999, Compositional strati-fi cation in the deep mantle: Science, v. 283, p. 1881–1884.

Klein, E.M., and Karsten, J.L., 1995, Ocean-ridge basalts with convergent-margin geochemical affi nities from the Chile Ridge: Nature, v. 374, no. 6517, p. 52–57, doi: 10.1038/374052a0.

Komor, S.C., Grove, T.L., and Hebert, R., 1990, Abyssal peridotites from ODP Hole 670a (21°10N, 45°02W): Residues of mantle melting exposed by non-constructive axial divergence, in Detrick, R., Honnorez, J., Bryan, W.B., and Juteau, T., eds., Proceedings of the Ocean Drilling Program: Scientifi c Results, Volume 106/109: College Station, Texas, Ocean Drill-ing Program, p. 85–101.

Konstantopoulou, G., and Economou-Eliopoulos, M., 1990, Geochemistry of the Vourinos chromite ores, Greece, in Malpas, J., Moores, E.M., Panayiotou, A., and Xenophontos, C., eds., Ophiolites: Oceanic Crustal Analogues: Proceedings of the Symposium “Troodos 1987”: Cyprus, Nicosia, Geological Survey Department, p. 605–613.

Laurent, R., 1992, Peridotite intrusions emplaced in the fossil suprasubduction zone environment of Cyprus, in Parson, L.M., Murton, B.J., and Brown-ing, P., eds., Ophiolites and Their Modern Oceanic Analogues: Geological Society [London] Special Publication 60, p. 233–239.

Lehnert, K., Su, Y., Langmuir, C.H., Sarbas, B., and Nohl, U., 2000, A global geochemical database structure for rocks: Geochemistry, Geophysics: Geosystems, v. 1, no. 5, doi: 10.1029/1999GC000026.

Le Mee, L., Girardeau, J., and Monnier, C., 2004, Mantle segmentation along the Oman ophiolite fossil mid-ocean ridge: Nature, v. 532, p. 167–172.

le Roux, P.J., le Roux, A.P., Schilling, J.-G., Shimizu, N., Perkins, W.W., and Pearce, N.J.G., 2002, Mantle heterogeneity beneath the southern Mid-Atlantic Ridge: Trace element evidence for contamination of ambient asthenospheric mantle: Earth and Planetary Science Letters, v. 203, p. 479–498, doi: 10.1016/S0012-821X(02)00832-4.

Liati, A., Gebauer, D., and Fanning, C.M., 2004, The age of ophiolitic rocks of the Hellenides (Vourinos, Pindos, Crete); fi rst ion probe (SHRIMP) zircon ages: Chemical Geology, v. 207, p. 171–188.

Lippard, S.J., Shelton, A.W., and Gass, I.G., 1986, The Ophiolite of Northern Oman: Geological Society of London Memoir 11, 178 p.

MacLeod, C.J., Allerton, S., Gass, I.G., and Xenophontos, C., 1990, Structure of a fossil ridge-transform intersection in the Troodos ophiolite: Nature, v. 348, p. 717–720.

MacPherson, G.J., Phipps, S.P., and Grossman, J.N., 1990, Diverse sources for igneous blocks in Franciscan melanges, California Coast Ranges: The Journal of Geology, v. 98, no. 6, p. 845–862.

Malpas, J., and Langdon, G., 1984, Petrology of the Upper Pillow Lava suite, Troodos ophiolite, Cyprus, in Gass, I.G., Lippard, S.J., and Shelton, A.W., eds., Ophiolites and Oceanic Lithosphere: Geological Society [London] Special Publication 60, p. 155–167.

Malpas, J., Xenophontos, C., and Williams, D., 1992, The Ayia Varvara Formation of SW Cyprus; a product of complex collisional tectonics: Tectonophysics, v. 212, no. 3–4, p. 193–211, doi: 10.1016/0040-1951(92)90291-D.

Malpas, J., Calon, T., and Squires, G., 1993, The development of a Late Cretaceous microplate suture zone in SW Cyprus, in Prichard, H.M., Alabaster, T., Harris, N.B.W., and Neary, C.R., eds., Magmatic Processes and Plate Tectonics: Geological Society [London] Special Publication 76, p. 177–195.

Maxeiner, R.O., Corrigan, D., Harper, C.T., MacDougall, D.G., and Ansdell, K., 2005, Paleoproterozoic arc and ophiolitic rocks of the northwest-margin of the Trans-Hudson orogen, Saskatchewan, Canada: Their contribution to a revised tectonic framework for the orogen: Precambrian Research, v. 136, p. 67–106.

McCulloch, M.T., and Cameron, W.E., 1983, Nd-Sr isotopic study of primi-tive lavas from the Troodos ophiolite, Cyprus; evidence for a subduction-related setting: Geology, v. 11, no. 12, p. 727–731, doi: 10.1130/0091-7613(1983)11<727:NISOPL>2.0.CO;2.

Metcalf, R.V., Wallin, E.T., Willse, K.R., and 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., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 403–418.

Miyashiro, A., 1973, The Troodos ophiolitic complex was probably formed in an island arc: Earth and Planetary Science Letters, v. 19, no. 2, p. 218–224, doi: 10.1016/0012-821X(73)90118-0.


Moores, E.M., 1998, Ophiolites, the Sierra Nevada, “Cordillera,” and orogeny along the Pacifi c and Caribbean margins of North and South America: International Geology Review, v. 40, no. 1, p. 40–54.

Moores, E.M., 2002, Pre–1 Ga (pre-Rodinian) ophiolites: Their tectonic and environmental implications: Geological Society of America Bulletin, v. 114, p. 80–95, doi: 10.1130/0016-7606(2002)114<0080:PGPROT>2.0.CO;2.

Moores, E.M., 2003, A personal history of the ophiolite concept, in Dilek, Y., and Newcomb, S., Ophiolite Concept and Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 17–29.

Moores, E.M., and Vine, F.J., 1971, Troodos massif, Cyprus, and other ophio-lites as oceanic crust: Evaluation and implications: Philosophical Trans-actions of the Royal Society of London, v. 268, p. 443–466.

Moores, E.M., Robinson, P.T., Malpas, J., and Xenophontos, C., 1984, Model for the origin of the Troodos Massif, Cyprus, and other Mideast ophio-lites: Geology, v. 12, no. 8, p. 500–503, doi: 10.1130/0091-7613(1984)12<500:MFTOOT>2.0.CO;2.

Moores, E.M., Kellogg, L.H., and Dilek, Y., 2000, Tethyan ophiolites, man-tle convection, and tectonic “historical contingency”; a resolution of the “ophiolite conundrum,” in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 3–12.

Muenow, D.W., Perfi t, M.R., and Aggrey, K.E., 1991, Abundances of volatiles and genetic relationships among submarine basalts from the Woodlark basin: Southwest Pacifi c: Geochimica et Cosmochimica Acta, v. 55, no. 8, p. 2231–2239, doi: 10.1016/0016-7037(91)90099-Q.

Murchey, B.L., 1984, Biostratigraphy and lithostratigraphy of chert in the Franciscan Complex, Marin Headlands, California, in Blake, C., Jr., ed., Franciscan Geology of Northern California: Los Angeles, Society of Economic Paleontologists and Mineralogists, Pacifi c Section, Field Trip Guidebook 43, p. 51–70.

Murchey, B.L., and Jones, D.L., 1984, Age and signifi cance of chert in the Franciscan Complex, in the San Franciscan Bay region, in Blake, C., Jr., ed., Franciscan Geology of Northern California: Los Angeles, Society of Economic Paleontologists and Mineralogists, Pacifi c Section, Field Trip Guidebook 43, p. 23–30.

Nehlig, P., 1993, Interactions between magma chambers and hydrothermal sys-tems; oceanic and ophiolitic constraints: Journal of Geophysical Research, ser. B, Solid Earth and Planets, v. 98, no. 11, p. 19,621–19,633.

Niu, Y., Collerson, K.D., Batiza, R., Wendt, J.I., and Regelous, M., 1999, Origin of enriched-type mid-ocean ridge basalt at ridges far from mantle plumes: The East Pacifi c Rise at 11°N: Journal of Geophysical Research, v. 104, no. B4, p. 7067–7087, doi: 10.1029/1998JB900037.

Ohara, Y., Stern, R.J., Ishii, T., Yurimoto, H., and Yamazaki, T., 2002, Peridotites from the Mariana Trough: First look at the mantle beneath an active back-arc basin: Contributions to Mineralogy and Petrology, v. 143, p. 1–18.

Okamura, H., Arai, S., and Kretschmar, L.M., 2006, Petrology of forearc peridotite from the Hahajima Seamount, the Izu-Bonin arc, with special reference to chemical characteristics of chromian spinel: Mineralogical Magazine, v. 70, p. 15–26, doi: 10.1180/0026461067010310.

O’Nions, R.K., Hamilton, P.J., and Evensen, N.M., 1977, Variation in 143Nd/144Nd and 87Sr/86Sr in oceanic basalts: Earth and Planetary Science Letters, v. 63, p. 229–240, doi: 10.1016/0012-821X(83)90039-0.

Parkinson, I.J., and Pearce, J.A., 1998, Peridotites from the Izu-Bonin-Mariana forearc (ODP Leg 125): Evidence for mantle melting and melt-mantle interaction in a supra-subduction zone setting: Journal of Petrology, v. 39, p. 1577–1618.

Pearce, J.A., 1982, Trace element characteristics of lavas from destructive plate boundaries, in Thorpe, R.S., ed., Andesites: Orogenic Andesites and Related Rocks: Chichester, UK, John Wiley & Sons, p. 525–548.

Pearce, J.A., 2003, Suprasubduction zone ophiolites: The search for modern analogues, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 269–294.

Pearce, J.A., and Parkinson, I.J., 1993, Trace element models for mantle melting; application to volcanic arc petrogenesis, in Prichard, H.M., Alabaster, T., Harris, N.B.W., and Neary, C.R., eds., Magmatic Processes and Plate Tectonics: Geological Society [London] Special Publication 76, p. 373–403.

Pearce, J.A., Lippard, S.J., and Roberts, S., 1984, Characteristics and tectonic signifi cance of supra-subduction zone ophiolites, in Kokelaar, B.P., and

Howells, M.F., eds., Marginal Basin Geology: Palo Alto, California, Blackwell Scientifi c Publications, p. 74–94.

Pearce, J.A., Baker, P.E., Harvey, P.K., and Luff, I.W., 1995, Geochemical evi-dence for subduction fl uxes, mantle melting and fractional crystallization beneath the South Sandwich Island arc: Journal of Petrology, v. 36, no. 4, p. 1073–1109.

Pearce, J.A., Stern, R.J., Bloomer, S.H., and Fryer, P., 2005, Geochemical map-ping of the Mariana arc-basin system: Implications for the nature and distribution of subduction components: Geochemistry, Geophysics, Geo-systems (G3), v. 6, no. 7, doi: 10.1029/2004GC000895,

Penrose Conference Participants, 1972, Penrose Field Conference on Ophio-lites: Geotimes, v. 17, p. 24–25.

Perfi t, M.R., Gust, D.A., Bence, A.E., Arculus, R.J., and Taylor, S.R., 1980, Chemical characteristics of island-arc basalts: Implications for mantle sources: Chemical Geology, v. 30, p. 227–256.

Perfi t, M.R., Langmuir, C.H., Baekisapa, M., Chappell, B.W., Johnson, R.W., Staudigel, H., and Taylor, S.R., 1987, Geochemistry and petrology of volcanic rocks from the Woodlark basin; addressing questions of ridge subduction, in Taylor, B., and Exon, N.F.E., eds., Marine Geology, Geo-physics and Geochemistry of the Woodlark Basin–Solomon Islands: Menlo Park, California, Circum-Pacifi c Council for Energy and Mineral Resources, Earth Science Series, p. 113–154.

Pessagno, E.A., Jr., Hull, D.M., and Hopson, C.A., 2000, Tectonostratigraphic signifi cance 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., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geo-logical Society of America Special Paper 349, p. 383–394.

Pyle, D.G., Christie, D.M., and Mahoney, J.J., 1992, Resolving an iso topic boundary within the Australian-Antarctica discordance: Earth and Plan-etary Science Letters, v. 112, p. 161–178, doi: 10.1016/0012-821X(92)90014-M.

Pyle, D.G., Christie, D.M., Mahoney, J.J., and Duncan, R.A., 1995, Geo-chemistry and geochronology of ancient southeast Indian and south-west Pacifi c seafl oor: Journal of Geophysical Research, v. 100, no. B11, p. 22,261–22,282, doi: 10.1029/95JB01424.

Rampone, E., and Piccardo, G.B., 2000, The ophiolite-oceanic lithosphere analogue; new insights from the northern Apennines (Italy), in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 21–34.

Rautenschlein, M., Jenner, G.A., Hertogen, J., Hofmann, A.W., Kerrich, R., Schmincke, H.U., and White, W.M., 1985, Isotopic and trace element composition of volcanic glasses from the Akaki Canyon, Cyprus; implica-tions for the origin of the Troodos ophiolite: Earth and Planetary Science Letters, v. 75, no. 4, p. 369–383, doi: 10.1016/0012-821X(85)90180-3.

Regelous, M., Nui, Y., Wendt, J.I., Batiza, R., Greig, A., and Collerson, K.D., 1999, Variations in the geochemistry of magmatism on the East Pacifi c Rise at 10°30N since 800 ka: Earth and Planetary Science Letters, v. 168, p. 45–63, doi: 10.1016/S0012-821X(99)00048-5.

Robertson, A.H.F., 1986, Geochemical evidence for the origin of Late Triassic mélange units in the Oman Mountains as a small ocean basin formed by continental rifting: Earth and Planetary Science Letters, v. 77, no. 3–4, p. 318–332, doi: 10.1016/0012-821X(86)90143-3.

Robertson, A., and Xenophontos, C., 1993, Development of concepts concern-ing the Troodos ophiolite and adjacent units in Cyprus, in Prichard, H.M., Alabaster, T., Harris, N.B.W., and Neary, C.R., eds., Magmatic Processes and Plate Tectonics: Geological Society [London] Special Publication 76, p. 85–119.

Robinson, P.T., and Malpas, J., 1990, The Troodos ophiolite: new perspective on its origin and emplacement, in Malpas, J., Moores, E.M., Panayiotou, A., and Xenophontos, C., eds., Ophiolites: Oceanic Crustal Analogues: Proceedings of the Symposium “Troodos 1987”: Cyprus, Nicosia, Geo-logical Survey Department, p. 13–26.

Robinson, P.T., Melson, W.G., O’Hearn, T., and Schminke, H.-U., 1983, Vol-canic glass compositions of the Troodos ophiolite, Cyprus: Geology, v. 11, p. 400–404, doi: 10.1130/0091-7613(1983)11<400:VGCOTT>2.0.CO;2.

Robinson, P.T., Malpas, J., and Xenophontas, C., 2003, The Troodos Massif of Cyprus: Its role in the evolution of the ophiolite concept, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 295–308.


Rochette, P., Jenatton, L., Dupuy, C., Boudier, F., and Reuber, I., 1991, Dia-base dikes emplacement in the Oman ophiolite; a magnetic fabric study with reference to geochemistry, in Peters, T., Nicolas, A., and Coleman, R.J., eds., Ophiolite Genesis and Evolution of the Oceanic Lithosphere: Muscat, Oman: Norwell, Massachusetts, Kluwer Academic Publisher, p. 55–82.

Rogers, R.D., Karason, H., and van der Hilst, R.D., 2002, Epeirogenic uplift above a detached slab in northern Central America: Geology, v. 30, p. 1031–1034, doi: 10.1130/0091-7613(2002)030<1031:EUAADS>2.0.CO;2.

Saccani, E., Padoa, E., and Photiades, A., 2003, Triassic mid-ocean ridge basalts from the Argolis Peninsula (Greece); new constraints for the early oceanization phases of the Neo-Tethyan Pindos Basin, in Dilek, Y., and Robinson Paul, T., eds., Ophiolites in Earth History: Geological Society [London] Special Publication 218, p. 109–128.

Saltus, R.W., Hudson, T.L., Karl, S.M., and Morin, R.L., 2003, Rooted Brooks Range ophiolite; implications for Cordilleran terranes: Reply: Geology, v. 31, no. 1, p. 92, doi: 10.1130/0091-7613(2003)031<0092:>2.0.CO;2.

Saunders, A.D., Tarney, J., Marsh, N.G., and Wood, D.A., 1980, Ophiolites as ocean crust or marginal basin crust; a geochemical approach, in Ophio-lites: Proceedings of the International Ophiolite Symposium: Cyprus, Nicosia, p. 193–204.

Schilling, J.-G., 1973, Iceland mantle plume: Geochemical study of Reykjanes Ridge: Nature, v. 242, p. 565–571, doi: 10.1038/242565a0.

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., and Xenophontos, C., eds., Ophiolites: Oceanic Crustal Analogues: Pro-ceedings of the Symposium “Troodos 1987”: Cyprus, Nicosia, Geological Survey Department, p. 507–520.

Shervais, J.W., 2001, Birth, death, and resurrection: The life cycle of supra-subduction zone ophiolites: Geochemistry, Geophysics, Geosystems, v. 2, doi: 10.1029/2000GC000080.

Shervais, J.W., 2006, Signifi cance of subduction-related accretionary com-plexes in early Earth processes, in Reimold, U., and Gibson, R., eds., Early Earth Processes: Geological Society of America Special Paper 405, p. 173–192.

Shervais, J.W., and Kimbrough, D.L., 1985, Alkaline and transalkaline meta-basalts in the Franciscan Complex mélange, California, in Morris, E.M., and Pasteris, J.D., eds., Mantle Metasomatism and Alkaline Magmatism: Geological Society of America Special Paper 215, p. 165–183.

Shervais, J.W., Kimbrough, D.L., Renne, P.R., Hanan, B.B., Murchey, B., Snow, C.A., Zoglman-Schuman, M.M., and 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, v. 46, no. 4, p. 289–315.

Shervais, J.W., Murchey, B., Kimbrough, D.L., Renne, P., and Hanan, B.B., 2005a, Radioisotopic and biostratigraphic age relations in the Coast Range ophiolite, northern California: Implications for the tectonic evolu-tion of the western Cordillera: Geological Society of America Bulletin, v. 117, no. 5/6, p. 633–653, doi: 10.1130/B25443.1.

Shervais, J.W., Zoglman-Schuman, M.M., and Hanan, B.B., 2005b, The Stony-ford volcanic complex: A forearc seamount in the northern California Coast Ranges: Journal of Petrology, v. 46, no. 10, p. 2091–2128, doi: 10.1093/petrology/egi048.

Sisson, T.W., and Grove, T.L., 1993, Experimental investigations of the role of H2O in calc-alkaline differentiation and subduction zone magmatism: Contributions to Mineralogy and Petrology, v. 113, p. 143–166, doi: 10.1007/BF00283225.

Sisson, V.B., Roeske, S.M., and Pavlis, T.L., eds., 2003, Geology of a Trans-pressional Orogen Developed during Ridge-Trench Interaction along the North Pacifi c Margin: Geological Society of America Special Paper 371, 375 p.

Staudigel, H., McCulloch, M.T., Zindler, A., and Perfi t, M.R., 1987, Complex ridge subduction and island arc magmatism; an isotopic study of the New Georgia forearc and the Woodlark basin, in Taylor, B., and Exon, N.F., eds., Marine Geology, Geophysics and Geochemistry of the Woodlark Basin–Solomon Islands: Menlo Park, California, Circum-Pacifi c Council for Energy and Mineral Resources, Earth Science Series, p. 227–240.

Steinmann, G., 1906, Die ophiolitischen Zonen in den Mediterranean Kettenge-birgen (translated by Bernoulli, D., and Friedman, G.M.), in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geological Thought: Geological Society of America Special Paper 373, p. 77–92.

Stern, C.R., 1980, Geochemistry of Chilean ophiolites; evidence for the compo-sitional evolution of the mantle source of back-arc basin basalts: Journal of Geophysical Research, v. 85, no. B2, p. 955–966.

Stern, R.J., 2004, Subduction initiation: Spontaneous and induced: Earth and Planetary Science Letters, v. 226, p. 275–292, doi: 10.1016/S0012-821X(04)00498-4.

Stern, R.J., and Bloomer, S.H., 1992, Subduction zone infancy; examples from the Eocene Izu-Bonin-Mariana and Jurassic California arcs: Geological Society of America Bulletin, v. 104, no. 12, p. 1621–1636, doi: 10.1130/0016-7606(1992)104<1621:SZIEFT>2.3.CO;2.

Stern, R.J., Morris, J.D., Bloomer, S.H., and Hawkins, J.W., Jr., 1991, The source of the subduction component in convergent margin magmas; trace element and radiogenic isotope evidence from Eocene boninites, Mariana forearc: Geochimica et Cosmochimica Acta, v. 55, no. 5, p. 1467–1481, doi: 10.1016/0016-7037(91)90321-U.

Stoffers, P., Worthington, T., Hakinian, R., Petersen, S., Hannington, M.D., Türkay, M., and Scientifi c Shipboard Party, 2002, Silicic volcanism and hydrothermal activity documented at Pacifi c-Antarctic Ridge: Eos (Trans-actions, American Geophysical Union), v. 83, no. 28, p. 303–304, doi: 10.1029/2002EO000215.

Sturm, M.E., Klein, E.M., Graham, D.W., and Karsten, J., 1999, Age con-straints on crustal recycling to the mantle beneath the southern Chile Ridge; He-Pb-Sr-Nd isotope systematics: Journal of Geophysical Research, ser. B, Solid Earth and Planets, v. 104, no. B3, p. 5097–5114, doi: 10.1029/1998JB900107.

Sturm, M.E., Klein, E.M., Karsten, J.L., and Karson, J.A., 2000, Evidence for subduction-related contamination of the mantle beneath the south-ern Chile Ridge; implications for ambiguous ophiolite compositions, in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 13–20.

Suhr, G., and Edwards, S.J., 2000, Contrasting mantle sequences exposed in the Lewis Hills Massif; evidence for the early arc-related history of the Bay of Islands ophiolite, in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 433–460.

Sun, S.-S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the Ocean Basins: Geological Society [London] Special Publication 42, p. 313–345.

Taylor, B., and Exon, N.F., 1987, An investigation of ridge subduction in the Woodlark-Solomons region: Introduction and overview, in Taylor, B., and Exon, N.F., eds., Marine Geology, Geophysics and Geochemistry of the Woodlark Basin–Solomon Islands: Menlo Park, California, Circum-Pacifi c Council for Energy and Mineral Resources, Earth Science Series, p. 1–24.

Taylor, R.N., Murton, B.J., and Nesbitt, R.W., 1992, Chemical transects across intra-oceanic arcs: Implications for the tectonic setting of ophiolites, in Parson, L.M., Murton, B.J., Browning, P., eds., Ophiolites and Other Modern Oceanic Analogues: Geological Society [London] Special Pub-lication 60, p. 117–132.

Tian, L., Zhao, G., Zhao, G., Shi, X., and Hailong, L., 2005, Geochemistry of basaltic lavas from the Mariana Trough: Evidence for infl uence of sub-duction component on generation of backarc basin magmas: International Geology Review, v. 47, p. 387–397.

Umino, S., and Kushiro, I., 1989, Experimental studies on boninite petro-genesis, in Crawford, A.J., ed., Boninites: London, UK, Unwin Hyman, p. 89–111.

Van der Laan, S.R., Flower, M.F.J., and Koster van Groos, A.F., 1989, Experi-mental evidence for the origin of boninites; near-liquidus phase rela-tions to 7.5 kbar, in Crawford, A.J., ed., Boninites: London, UK, Unwin Hyman, p. 112–147.

Van Der Laan, S.R., Arculus, R.J., Pearce, J.A., and Murton, B.J., 1992, Petrography, mineral chemistry and phase relations of the basement boninite series of Site 786, Izu-Bonin forearc, in Fryer, P., Pearce, J.A., and Strokking, L.B., eds., Proceedings of the Ocean Drilling Program Scientifi c Results, Volume 125: College Station, Texas, Ocean Drilling Program, p. 171–201.

Varne, R., Brown, A.V., and Falloon, T., 2000, Macquarie Island: Its geology, structural history, and the timing and tectonic setting of its N-MORB to E-MORB magmatism, in Dilek, Y., Moores, E.M., Elthon, D., and


Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 301–320.

Volker, F., Alther, R., Jochum, K.-P., and McCulloch, M.T., 1997, Quaternary volcanic activity of the southern Red Sea: New data and assessment of models on magma sources and Afar plume-lithosphere interaction: Tec-tonophysics, v. 278, p. 15–29, doi: 10.1016/S0040-1951(97)00092-9.

Walker, D.A., and Cameron, W.E., 1983, Boninite primary magmas; evidence from the Cape Vogel Peninsula: PNG: Contributions to Mineralogy and Petrology, v. 83, no. 1–2, p. 150–158, doi: 10.1007/BF00373088.

Wallin, E.T., and Metcalf, R.V., 1998, Supra-subduction zone ophiolite formed in an extensional forearc: Trinity terrane, Klamath Mountains, California: The Journal of Geology, v. 106, p. 591–608.

Weaver, B.L., 1991, The origin of ocean island basalt end-member compositions: Trace element and isotopic constraints: Geology, v. 104, p. 381–397.

Weins, D.A., and Smith, G.P., 2003, Seismological constraints on structure and fl ow patterns within the mantle wedge, in Eiler, J., ed., Inside the

Subduction Factory: Active Margins and Marginal Basins of the West-ern Pacifi c: American Geophysical Union Geophysical Monograph 138, p. 59–82.

Wood, D.A., 1980, The application of a Th-Hf-Ta diagram to problems of tectonomagmatic classifi cation and to establishing the nature of crustal contamination of basaltic lavas of the British Tertiary volcanic prov-ince: Earth and Planetary Science Letters, v. 50, no. 1, p. 11–30, doi: 10.1016/0012-821X(80)90116-8.

Woodhead, J.D., Eggins, S.M., and Johnson, R.W., 1998, Magma genesis in the New Britain island arc: Further insights into melting and mass trans-fer processes: Journal of Petrology, v. 39, p. 1641–1668, doi: 10.1093/petrology/39.9.1641.

Zindler, A., and Hart, S., 1986, Chemical geodynamics: Annual Review of Earth and Planetary Sciences, v. 14, p. 493–571, doi: 10.1146/annurev.ea.14.050186.002425.

MANUSCRIPT ACCEPTED BY THE SOCIETY 16 JULY 2007

Printed in the USA

suprasubduction-zone ophiolites: is there really an ... - my articles/gsa-sp438-2008... ·...

Documents