Ophiolite Genesis and Global Tectonics -Dilek & Furnes 2011

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Geological Society of America Bulletin Ophiolite genesis and global tectonics: Geochemical and tectonic fingerprinting of ancient oceanic lithosphere Yildirim Dilek and Harald Furnes Geological Society of America Bulletin 2011;123;387-411 doi: 10.1130/B30446.1

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I NVITED R EVIEW A RTICLE

Ophiolite genesis and global tectonics: Geochemical and tectonic fingerprinting of ancient oceanic lithosphere Yildirim Dilek1,† and Harald Furnes 2 1 Department of

Geology,, Shideler Hall, Miami University Geology University,, Oxford, Oxford, Ohio 45056, USA, and Faculty Faculty of Earth Sciences, Sciences, China University of Geosciences at Wuhan, Wuhan 430074, Hubei Province, China 2 Department of Earth Science & Centre Centre for Geobiology, Geobiology, University of Bergen, Bergen, Bergen 5007, Norway Norway

ABSTRACT Ophiolites, and discussions on their origin and significance in Earth’s history, have been instrumental in the formulation, testing, and establishment of hypotheses and theories in earth sciences. The definition, tectonic origin, and emplacement mechanisms of ophiolites have been the subject of a dynamic and continually evolving concept since the nineteenth century. Here, we present a review of these ideas as well as a new classification of ophiolites, incorporating the diversity in their structural architecture and geochemical signatures that results from variations in petrological, geochemical, and tectonic processes during formation in different geodynamic settings. We define ophiolites as suites of temporally and spatially associated ultramafic to felsic rocks related to separate melting episodes and processes of magmatic differentiation in particular tectonic environments. Their geochemical characteristics, internal structure, and thickness vary with spreading rate, proximity to plumes or trenches, mantle temperature, mantle fertility, and the availability of fluids. Subductionrelated ophiolites include suprasubductionzone and volcanic-arc types, the evolution of which is governed by slab dehydration and accompanying metasomatism of the mantle, melting of the subducting sediments, and repeated episodes of partial melting of metasomatized peridotites. Subduction-unrelated ophiolites include continental-margin, midocean-ridge (plume-proximal, plume-distal, and trench-distal), and plume-type (plumeproximal ridge and oceanic plateau) ophio-

E-mail: dileky@mu [email protected]. ohio.edu.

†

GSA Bulletin; March/April 2011; v. 123; no.

lites that generally have mid-ocean-ridge basalt (MORB) compositions. Subductionrelated lithosphere and ophiolites develop during the closure of ocean basins, whereas subduction-unrelated types evolve during rift drift and seafloor spreading. The peak times of ophiolite genesis and emplacement in Earth history coincided with collisional events leading to the construction of supercontinents, continental breakup, and plumerelated supermagmatic events. Geochemical and tectonic fingerprinting of Phanerozoic ophiolites within the framework of this new ophiolite classification is an effective tool for identification of the geodynamic settings of oceanic crust formation in Earth history, and it can be extended into Precambrian greenstone belts in order to investigate the ways in which oceanic crust formed in the Archean. INTRODUCTION

Ophiolites represent fragments of upper mantle and oceanic crust (Dewey and Bird, 1971; Coleman, 1977; Nicolas, 1989) that were incorporated into continental margins during continent-continent and arc-continent collisions (Dilek and Flower, 2003), ridge-trench interactions (Cloos, 1993; Lagabrielle et al., 2000), and/or subduction-accretion events (Cawood et al., 2009). They are generally found along suture zones in both collisional-type (i.e., Alpine, Himalayan, Appalachian) and accretionary-type (i.e., North American Cordilleran) orogenic belts (Fig. 1) that mark major boundaries between amalgamated plates or accreted terranes (Lister and Forster, 2009). The geological record of the evolution of ocean basins from the rift-drift and seafloor spreading stages to the initiation of subduction and final closure (the Wil-

son cycle) is well preserved in most orogenic belts. Magmatism during each of these phases produces spatially and temporally associated, mafic-ultramafic to highly evolved rock assemblages. These rock units, which have varying internal structures, geochemical affinities, and age ranges, and originally formed in different geodynamic settings, constitute discrete ophiolite complexes and can become tectonically juxtaposed in collision zones (Dilek, 2003). In the Penrose definition (Anonymous, 1972, p. 24), an ophiolite is described as a “distinctive assemblage of mafic to ultramafic rocks” that includes, from bottom to top, tectonized peridotites, cumulate peridotites, and pyroxenites overlain by layered gabbros, sheeted basaltic dikes, a volcanic sequence, and a sedimentary cover; an ophiolite may be incomplete, tectonically dismembered, or metamorphosed. This original Penrose definition of ophiolites (Anonymous, 1972) is highly restrictive and does not reflect the actual heterogeneity in ophiolite composition and occurrence, and therefore a more deterministic approach to defining ophiolites and their igneous evolution is needed. In this paper, we first review the evolution of the ophiolite concept before and after the formal Penrose definition, and we redefine an ophiolite in light of recent observations and diverse data sets from ophiolites worldwide. We outline the significance of ophiolite pulses in Earth history within a global tectonic framework and introduce a new and more comprehensive classification of ophiolites based on their distinctive internal structures, geochemical signatures, and regional tectonics. We then present petrogenetic models for the formation of different types of ophiolites and discuss the implications of this new ophiolite classification for the origin of Precambrian oceanic crust, particularly for some Archean greenstone belts.

3/4; p. 387–411; doi: 10.1130/B30446.1; 12 figures; 2 tables, Data Repository item 2011131. For permission to copy, contact [email protected] © 2011 Geological Society of America

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mid-ocean-ridge

Indonesian belt (Cenozoic) Western Pacific and Cordilleran belts (Paleozoic-Tertiary) Alpine - Himalayan belt (Jurassic - Cretaceous) 7 5 °

15°W

7 5 ° E

Appalachian - Caledonian Hercynian - Uralian & Central Asian belts (early Paleozoic) Tasmanides (Paleozoic)

165°E

c h

S unda T r e n

Figure 1. Global distribution of major Phanerozoic orogenic belts and ophiolite age clusters on a north polar projection. Significant examples of different ophiolite types with characteristic geochemistries are mar ked with symbols used in Figure 2. Modern mid-ocean ridges and subduction zones (marked by trenches) where contemporary oceanic lithosphere has been produced are also depicted. The two major arc-trench rollback systems, Izu-Bonin-Mariana and Tonga-Kermadec, are the sites of ophiolite and volcanic-arc generation, which undergo tectonic extension and trenchward-migrating magmatic construction. The collision zone between the NW Australian passive margin and the Sunda arc-trench system where the island of Timor has been emerging during the last ~5 m.y. represents the best modern analogue for ophiolite emplacement.

HISTORICAL BACKGROUND AND NEW DEFINITION OF OPHIOLITES Early Ideas and Evolving Ophiolite Concept

The term “ophiolite” was first used in 1813 by a French mineralogist, Alexandre Brongniart (1770–1847), in reference to serpentinites in mélanges; he subsequently redefined his defini388

tion of an ophiolite (Brongniart,1821) to include a suite of magmatic rocks (ultramafic rocks, gabbro, diabase, and volcanic rocks) occurring in the Apennines. Gustav Steinmann (1856– 1929) elevated the “ophiolite” term to a new concept by defining ophiolites as spatially associated kindred rocks that originally formed as in situ intrusions in axial parts of geosynclines (Steinmann, 1927). Steinmann emphasized the common occurrence of peridotite (serpenti-

nite), gabbro, and diabase-spilite, in association with deep-sea sedimentary rocks in the Mediterranean mountain chains and interpreted the origin of these rocks as differentiated magmatic units evolved on the ocean floor. He considered these rock assemblages to have developed from a consanguineous igneous process during the evolution of eugeosynclines. This interpretation subsequently led to the widely known notion of the “Steinmann trinity.”

Geological Society of America Bulletin, March/April 2011

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Although Steinmann considered peridotite, gabbro, diabase, and volcanic rocks in ophiolites as comagmatic in origin, his observation that gabbroic and diabasic rocks were intrusive bodies in the serpentinized peridotites is an extremely important one because it differs from the contemporary interpretation of the “Penrose-type” ophiolite. It implies that, at least in the Apennine ophiolites, the gabbros and volcanic rocks are younger than the peridotites. Steinmann also correctly interpreted the ophiolites in the Northern Apennines as thrust sheets tectonically overlying the Tertiary sedimentary rocks in Tuscany (Steinmann, 1913). This interpretation led to the discovery of allochthonous nappe sequences in the Alpine-Apennine orogenic system. Thayer (1967) discussed the significance of the consanguineous relationship between ultramafic and associated mafic rocks in alpine-type peridotites, which were defined by Benson (1926) earlier, and explained how the gabbro, diabase, and other leucocratic rocks in alpine-type peridotites could have originated from a single primary peridotitic magma. Jackson and Thayer (1972) subsequently distinguished harzburgite-type versus lherzolitetype alpine peridotites. In this subgrouping, the harzburgite-type alpine peridotites represent the uppermost oceanic mantle, whereas the less-depleted lherzolite-type alpine peridotites correspond to the subcontinental mantle and/or to the deeper oceanic mantle, where partial melting is much less intense. Recent studies of ophiolites have shown that both harzburgiteand lherzolite-type peridotites may occur in ophiolites, and that they can be used to classify ophiolite types and their inferred spreading rates of formation in an oceanic setting (Ishiwatari, 1985; Boudier and Nicolas, 1985; Nicolas and Boudier, 2003). In his classic paper published in Crust of the Earth (Geological Society of America Special Paper 62), Hess (1955, p. 393) stated that Steinmann’s ophiolite concept was confusing because “it obscured critical relationships of its [ophiolite] various members to the tectonic cycle.” Recognizing the importance of serpentinites and alpine-type peridotites in orogeny and mountain-building episodes, he argued that serpentinites and rocks of Steinmann’s trinity are common in island arcs and that “island arcs represent an early stage in the development of an alpine-type of mountain system” (p. 395). Hess was, therefore, advocating an island-arc origin of mafic-ultramafic rock assemblages and serpentinized peridotites found in orogenic belts. This was nearly 20 yr before Miyashiro (1973) made the first formal and rather controversial call on the island-arc origin of the Troodos

ophiolite (Cyprus), connecting ophiolite genesis to subduction-zone processes. Hess discussed in his 1962 paper that the main oceanic crustal layer (his layer 3) along the Mid-Atlantic Ridge was made largely of serpentinite (his Fig. 2, p. 603; Hess 1962), and that the seismic velocity of this layer would be highly variable, depending on the magnitude of serpentinization of the peridotite. He proposed that the interface between the oceanic crust (composed mainly of serpentinite) and the underlying peridotite with seismic velocities of 7.4 km/s represented the Moho discontinuity. Since he had interpreted serpentinites as hydrated peridotites, Hess described the Moho beneath the Mid-Atlantic Ridge as an alteration front (phase transition) rather than a sharp boundary separating the igneous crust from the underlying mantle (his Fig. 7, p. 612). Although we now know that oceanic crust is not made of 70% serpentinite, marine geological and geophysical studies have documented that the slowspreading oceanic crust along the Mid-Atlantic Ridge has a highly heterogeneous lithological composition and thickness (Dick, 1989). For example, thin-crust domains along the ridge axis (i.e., magma-poor segment ends) consist of tectonically uplifted ultramafic rocks with gabbroic intrusions and a thin basaltic cover (Cannat et al., 1995). This nonuniform thickness and the heterogeneous lithostratigraphy of the MidAtlantic Ridge crust are remarkably similar to Steinmann’s description of the Ligurian ophiolites in the Apennines. It also largely corresponds to Hess’ characterization of oceanic crust developed at the Mid-Atlantic Ridge. This “Hess-type crust” differs significantly from “Penrose-type” oceanic crust in terms of its internal architecture, as discussed in the following. The Dutch geologist de Roever (1957) reinterpreted the Steinmann trinity to result of mantle melting, producing the basaltic rocks on top and the residual ultramafic rocks at the bottom. Subsequently, the Swiss petrologist Vuagnat argued that the peridotite massifs in ophiolites were partial melting residues in the upper mantle (Vuagnat, 1964), because he thought that the overwhelming abundance of ultramafic rocks in ophiolites compared to the small volumetric occurrence of gabbroic rocks could not simply be explained by differentiation of submarine outpourings of basaltic magma. It is important to note that these two papers by de Roever (1957) and (Vuagnat, 1964) mark in the literature the beginning of a significant shift in Steinmann’s “cogenetic” ophiolite concept and of a new paradigm in oceanic crustal evolution. Recognition of extensional sheeted dike complexes, the existence of a refractory mantle unit represented by harzburgitic peridotites

with high-temperature deformation fabrics, fossil magma chambers in plutonic sequences, and the allochthonous nature of ophiolites by the mid-1960s was instrumental in the formulation of the ophiolite model and the ophiolite– ocean crust analogy within the framework of the new plate-tectonic theory. The ophiolite suite became an ideal analogue to explain the seismic velocity structure of modern oceanic lithosphere, as more seismic data became available from modern ocean basins, particularly from the Pacific Ocean. Combined with observations from the Troodos (Cyprus) and Semail (Oman) ophiolites in particular, the seismic velocity structure of modern oceanic crust and its inferred layer-cake pseudostratigraphy came to be known as the “ophiolite model.” This analogy was confirmed at the first Penrose Conference on ophiolites in 1972 (Anonymous, 1972), whereby an ideal ophiolite sequence was defined to have a layer-cake pseudostratigraphy complete with a sheeted dike complex as a result of seafloor spreading. Ophiolites were interpreted to have developed mainly at ancient mid-ocean ridges through this model. In a uniformitarian approach, ophiolite geologists then started reconstructing the evolution of fossil oceanic lithosphere exposed on land as a product of paleo–mid-ocean ridges using the ophiolite–ocean crust analogy (Gass, 1968; Coleman, 1971; Moores and Vine, 1971; Cann, 2003, and references therein). Geochemical studies challenged this view of a mid-ocean-ridge origin of ophiolites as early as the beginning of the 1970s, and suggested the association of magma evolution with subduction zones. Miyashiro (1973, p. 218) argued that “about one-third of the analyzed rocks of the lower pillow lavas and sheeted dike rocks in the Troodos ophiolite follows a calc-alkalic trend,” suggesting that “the massif was created as a basaltic volcano in an island arc with a relatively thin ocean-type crust rather than in a mid-oceanic ridge.” This was the first formal proposal of a subduction-zone origin of the Troodos “oceanic crust” that questioned the “ruling hypothesis” of a mid-ocean-ridge setting of ophiolite genesis. Miyashiro’s geochemical argument on the island-arc origin of the Troodos ophiolite would start a major paradigm shift in the ophiolite concept in the wake of the platetectonic revolution. The subsequent scientific exchange in the form of discussions and replies to Miyashiro’s 1973 paper initiated a long-lasting debate about the tectonic setting of ophiolite genesis. Pearce (1975) proposed a marginal basin origin for the Troodos massif during the evolution of an incipient submarine island arc. Findings from modern subduction-zone environments in the western Pacific prompted


389


researchers to consider more rigorously the evolution of ophiolites in spreading environments within the upper plate of subduction zones (Hawkins, 1977, 2003; Pearce, 2003). This development, which came about as a collective result of ophiolite studies on land and marine geological and geophysical investigations in modern convergent margin settings in the oceans, led to the definition of suprasubduction-zone ophiolites in the early 1980s (Pearce et al., 1984). The forearc environment of the Izu-Bonin-Mariana arc-trench system is today one of the best studied (through deepocean drilling and submersible diving surveys) and best understood modern suprasubduction zones that we consider to be a contemporary suprasubduction-zone ophiolite factory (Fig. 1; Stern et al., 1989; Stern and Bloomer, 1992; Reagan et al., 2010; Dilek and Furnes, 2010). Systematic petrological and geochemical investigations of world ophiolites throughout the 1980s and 1990s demonstrated the significance of subduction-zone–derived fluids and melting history in development of ophiolitic magmas (Saunders and Tarney, 1984; Rautenschlein et al., 1985; Hébert and Laurent, 1990; Thy and Xenophontos, 1991; Beccaluva et al., 1994; Bédard et al., 1998; Dilek et al., 1999; Shervais, 2000; Dilek and Flower, 2003). Forearc, embryonic arc, and backarc settings in suprasubduction zones became the most widely accepted tectonic environments of origin. New Definition of Ophiolites

The basic tenet of the 1972 Penrose definition is that an ideal ophiolite has a layer-cake pseudostratigraphy with laterally persistent and horizontal contacts. The Mohorovicic discontinuity (Moho) is considered to be a petrological transition zone separating the crustal and upper-mantle rocks that have a melt-residua genetic relationship. Studies since 1972 have demonstrated, however, that most ophiolites have a dynamic evolution and display a laterally discontinuous and vertically heterogeneous crustal architecture and varying geochemical characteristics due to multiple magmatic episodes and different mantle sources during their igneous evolution. The fossil Moho also differs in character in ophiolites; in some, it represents a major tectonic discontinuity (i.e., detachment fault), whereas in some others, it is an alteration front. However, in some ophiolites it is a nearly 1-km-thick transition zone reminiscent of the Moho in slow-spreading young oceanic lithosphere (Dick et al., 2006). The diversity in the architecture and geochemical fingerprints observed in ophiolites reflects differences in igneous and tectonic processes involved in the 390

formation of oceanic crust in different geodynamic settings. We define an ophiolite as an allochthonous fragment of upper-mantle and oceanic crustal rocks that is tectonically displaced from its primary igneous origin of formation as a result of plate convergence. Such a slice should include a suite of, from bottom to top, peridotites and ultramafic to felsic crustal intrusive and volcanic rocks (with or without sheeted dikes) that can be geochronologically and petrogenetically related; some of these units may be missing in incomplete ophiolites. Ophiolite emplacement is a process that starts with displacement of oceanic lithosphere from its primary geodynamic environment and ends with its incorporation into mountain belts during orogenesis (Coleman, 1971; Dewey, 1976; Searle and Cox, 1999; Gray et al., 2000; Wakabayashi and Dilek, 2003). Ophiolites are commonly emplaced on a passive continental margin (buoyant crust) and island arc or in an accretionary complex. The magmatic and structural architecture of an ophiolite may reflect a product and complex interplay of successive melting episodes and processes of magmatic differentiation, spreading rate and geometry, intra-oceanic faulting, and deformation associated with tectonic extension, proximity to plumes or trenches, mantle temperature and fertility, and the availability of fluids during its primary igneous evolution. Some ophiolites are stratigraphically overlain by pelagic (chert or limestone) and/or Fe-Mn–rich hydrothermal sedimentary rocks and are underlain by amphibolite-greenschist rocks related to their tectonic displacement and emplacement. OPHIOLITE PULSES AND GLOBAL TECTONICS

The distribution of ophiolites in orogenic belts shows spatial and temporal patterns (Fig. 1), and the clusters of ophiolites with particular age ranges in different orogenic belts mark clear pulses, reflecting peak times of ophiolite genesis and emplacement in Earth history (Fig. 2). Some of the main ophiolite pulses overlap in time with major orogenic events that led to the construction of supercontinents. Examples include the Famatinian (Fmt) and Caledonian (Cld; Baltica- Laurentia collision) orogens in the early Paleozoic, which collectively formed the Gondwana and Laurasia supercontinents, and the AppalachianHercynian (Ap-Hy) and Altaid-Uralian (Al-Ur) orogens later in the Paleozoic, which built the Pangean supercontinent (Fig. 2; Moores et al., 2000). The sequential collisions of India (In-Eu) and Arabia (Ar-Eu) with Eurasia during the Neogene, after the emplacement of Neotethyan ophiolites and elimination of the Neotethyan sea-

ways by subduction, are part of the current assembly of a new supercontinent that has been taking place since the Paleogene. Paleozoic ophiolites in the AppalachianCaledonian orogenic belts (Fig. 1) developed in the Iapetus Ocean and its seaways between North America and Baltica-Avalonia (van Staal et al., 2009, and references therein). Ophiolites in Iberia, central Europe, and northwestern Africa evolved in the Rheic Ocean between BalticaAvalonia and Gondwana continental masses (Nance et al., 2010; Murphy et al., 2010, and references therein). The Paleozoic ophiolites in the Uralides and the Altaids in central Asia are the remnants of the Pleionic Ocean, which evolved between the Baltica–Eastern Europe and Kazakhstan-Siberian continental masses (Brown et al., 2006; Windley et al., 2002; Xiao et al., 2004). The Jurassic–Cretaceous ophiolites of the Tethyan Ocean systems extend from the BeticRif and Pyrenees in the west through the AlpineHimalayan orogenic belts in the center to the Indonesian region in the east (Fig. 1; Hall, 1997; Pubellier et al., 2004; Bortolotti and Principi, 2005). The Phanerozoic ophiolites in these collisional orogenic belts (i.e., Appalachian, Caledonides, Uralides, and Altaids in central Asia, Betic-Rif and Pyrenees, Alpine-Himalayan) commonly show mid-ocean-ridge basalt (MORB) to island-arc tholeiite (IAT) and boninitic geochemical affinities (Varfalvy et al., 1997; Bédard et al., 1998; Spadea and D’Antonio, 2006; Pagé et al., 2009). The ophiolites in the accretionarytype Western Pacific and Cordilleran orogenic belts are slivers of abyssal peridotites and volcanic ocean islands, seamounts, and mid-ocean-ridge crust scraped off from downgoing plates, and they are commonly associated with accretionary mélanges and high-pressure metamorphic rocks (Cloos, 1982; Wakabayashi, 1999; Ernst, 2005; Ring, 2008; Hall, 2009; Cawood et al., 2009; Xiao et al., 2010). The principal ophiolite pulses during the last 250 m.y. coincide with the emplacement of plume-related large igneous provinces (LIPs) and giant dike swarms (Ernst et al., 1995; Yale and Carpenter, 1998; Coffin and Eldholm, 2001) and collectively mark supermagmatic events in Earth history (Fig. 2). The enhanced large igneous province formation and ophiolite generation in the Late Jurassic and Cretaceous are particularly noteworthy (Vaughan and Scarrow, 2003). The evolution of the Tethyan and Caribbean ophiolites overlapped with the Cretaceous “superplume” event (120–80 Ma), which was responsible for the formation of oceanic plateaus in the Pacific and Indian Oceans, high global sea levels, and increased rates of seafloor spreading (Larson, 1991). The Jurassic–Cretaceous periCaribbean ophiolites (Fig. 1) include remnants



s e t i l o i h p o r o j a m f o r e b m u N

Central Asian ophiolites

Age (Ma) Ng

Pg

Tertiary

Cretaceous

Jurassic

Triassic

Permian

Mesozoic

Carb.

Devonian

Sil.

Ord.

Camb.

Paleozoic

Age (Ma)

h d c t e i h l a w e o r t e s r t a n s e e v t e i l o r i o j h a p M o

Figure 2. Ophiolite pulses and the distribution of major orogenic belts with ophiolite occurrences during the Phanerozoic. A. Ophiolite pulses and the geographic distribution of Phanerozoic ophiolites through time. B. Distribution of representative examples of major ophiolite types through time. C. Approximate time intervals for the lifespan of major supercontinents and their breakup, significant orogenic events, and supermagmatic events represented by the emplacement of giant dike swarms and large igneous provinces (LIPs). The main pulses of ophiolite generation coincide with plate movements leading to the closure of ocean basins and continental collisions, large magmatic events (with the production of large igneous provinces and giant dike swarms), and the breakup of supercontinents. Major orogenic events are (from youngest to oldest): Ar-Eu—Arabia-Eurasia collision, In-Eu—India-Eurasia collision, Al-Ur—Altaid-Uralian orogenies of Central Asia, Ap-Hy—Appalachian-Hercynian orogenies, Cld—Caledonian orogeny, Fmt—Famatinian orogeny, P-Af-Br—Pan-African– Brasiliano orogenies. Ng—Neogene; Pg—Paleogene. For a list of different ophiolite types, see Table 1.


391


of proto-Caribbean oceanic crust and the Caribbean–Colombian oceanic plateau (Kerr et al., 1998) and display a complex record of igneous activity associated with continental rifting, seafloor spreading, the construction of an oceanic plateau, and the development of island arcs (Giunta and Oliveri, 2009; Kerr et al., 2009). The most prominent ophiolite pulse during the Mesozoic coincided with the breakup of Pangea through discrete episodes of continental rifting during the Late Triassic and Jurassic (Fig. 2; Dalziel et al., 2000).

trolled the development of different ophiolite types in different tectonic environments (Dilek, 2003). We list representative examples of the main ophiolite types, their ages, geographic location, and related references in Table 1. These ophiolite types are marked in Figures 1 and 2 with different symbols, indicating formation in different tectonic environments, as explained in the following section. In Table 2, we also list and explain a series of abbreviations in reference to different ophiolite types and all the relevant geochemical terminology used in the next two sections and on the figures.

A NEW CLASSIFICATION OF OPHIOLITES

Tectonic Settings of Ophiolite Types

The main ophiolite pulses appear to be temporally and spatially linked to some first-order global tectonic and magmatic events. These global events and related mantle processes con-

Continental margin (CM) ophiolites form during the early stages of ocean basin evolution, following initial continental breakup. These ophiolites are fragments of magma-poor, ocean-

continent transitions (OCT). Modern, in situ ocean-continent transitions include the Iberia and Red Sea–Western Arabia rifted margins (Fig. 1). Some classic examples of continental margin ophiolites include the Jurassic ophiolites in the Northern Apennines (Ligurian) and the western Alps (Caby, 1995; Rampone et al., 2005; Manatschal and Müntener, 2009). These ophiolites consist of exhumed, subcontinental lithospheric mantle lherzolite directly overlain by basaltic lavas and intruded by small gabbroic plutons and rare mafic dikes. The crustal rocks display normal (N) MORB geochemical signatures. Continental margin ophiolites correspond to the lherzolite-type (LOT) ophiolites of Ishiwatari (1985) and Boudier and Nicolas (1985) and are the products of low degrees of melting of less-depleted subcontinental lithospheric mantle and upwelling asthenosphere (Rampone et al., 2005).

TABLE 1. REPRESENTATIVE EXAMPLES OF MAIN OPHIOLITE TYPES, THEIR GEOGRAPHIC LOCATIONS, APPROXIMATE AGES, AND RELATED REFERENCES Ophiolite Continental margin type 1 Tihama 2 Ligurian

Location

Age (Ma)

Red Sea, Saudi Arabia Italy

20 200

3 4 5

Ust-Belaya 1 Ust-Belaya 2 Nurali

NE Russia NE Russia S Urals, Russia

310 320 410

Mid-ocean-ridge type 1A 1B

Macquarie Isl. Taitao

SW Pacific S Chile

2

Khoy

Iran

4 5 6 7 8 9

Masirah Horo Kanai Kuyul 1 Kuyul 2 Kuyul 3 Nurali

W Indian Ocean Central Hokkaido, Japan NE Russia NE Russia NE Russia S Urals

Plume type 1A 1B 2 3 4 5 6 7 8

Loma de Hiero Bolivar Nicoya Peri-Caribbean 1 Peri-Caribbean 2 Duarte Loma La Monja Mino-Tamba 1 Mino-Tamba 2

Venezuela SW Colombia Costa Rica Cuba, Puerto Rica, Hispaniola Cuba, Puerto Rica, Hispaniola Hispaniola Hispaniola SW Japan SW Japan

80 80 89–95 105 125 140 155 185 200

Giunta et al. (2002) Nivia (1996) Kerr et al. (1997a, 1997b), Sinton et al. (1997), Hauff et al. (2000) Kerr et al. (1997a, 1997b), Giunta et al. (2006) Kerr et al. (1997a, 1997b), Giunta et al. (2006) Lapierre et al. (1997, 1999), Giunta et al. (2006), Escuder Viruete et al. (2009) Escuder Viruete et al. (2009) Ichiyama et al. (2008) Ichiyama et al. (2008)

Suprasubduction-zone type 1 Zambales 2 Antique 3A Troodos 3B Semail

Philippines Panay, Philippines Cyprus Oman

40–44 75–80 92–94 92–95

3C

Kizildag

Turkey

92–94

4 5 6A

Xigaze Sabah Mirdita

Tibet, China Northern Borneo Albania

120–126 135–140 160

6B 7 8

Pindos Cape Povorotny Yakuno

Greece Far East Asia SW Japan

160 230–250 270–280

Yumul et al. (2000), Encarnacion (2004) Dimalanta et al. (2006) Batanova and Sobolev (2000), Dilek and Furnes (2009) Lippard et al. (1986), Hacker et al. (1996), Warren et al. (2005) Dilek and Furnes (2009), Alabaster et al. (1982) Tinkler et al. (1981), Erendil (1984), Bagci et al. (2005), Dilek et al. (1999) Dilek and Thy (1998, 2009) Aitchison et al. (2003), Malpas et al. (2003), Zhang et al. (2003) Rangin et al. (1990), Müller (1991) Beccaluva et al. (1994), Bortolotti et al. (2002), Saccani and Photiades (2005) Dilek et al. (2007, 2008) Capedri et al. (1980), Saccani and Photiades (2005), Dilek and Furnes (2009) Sokolov et al. (2003) Ishiwatari (1985), Ichiyama and Ishiwatari (2004)

10 10 98-103 150 165–180 190 200 210 405

References Coleman et al. (1972, 1977), Dilek et al. (2009) Rampone and Piccardo (2000), Muntener and Piccardo (2003) Manatschale and Muntener (2009) Ishiwatari et at. (2003), Sokolov et al. (2003) Ishiwatari et at. (2003), Sokolov et al. (2003) Spadea et al. (2003) Kamentsky et al. (2000), Varne et al. (2000), Rivizzigno and Karson (2004) Le Moigne et al. (1996), Guivel et al. (1999), Lagabrielle et al. (2000), Shibuya et al. (2007) Ghazi and Hassanipak (2000), Hassanipak and Ghazi (2000) Khalatbari-Jafari et al. (2004) Peters and Mercolli (1998), Peters (2000) Ishiwatari et al. (2003) Sokolov et al. (2003) Sokolov et al. (2003) Sokolov et al. (2003) Pertsev et al. (1997), Spadea et al. (2003)

(continued )

392



Mid-ocean-ridge (MOR) ophiolites may form at plume-proximal (e.g., Iceland) and plumedistal mid-ocean ridges, trench-proximal midocean ridges, or trench-distal backarc spreading ridges (Table 2). They generally have a Penrosetype structural architecture (particularly at the centers of ridge segments) and show N-MORB (e.g., Argolis-Pindos in Greece), enriched (E) MORB (e.g., Macquarie Island), and/or contaminated (C) MORB geochemical affinities. N-MORB and E-MORB ophiolites have compositions that are more depleted and more enriched, respectively, than primitive mantle– derived magmas (Pearce, 2008). C-MORB ophiolites are crustally contaminated. The Taitao ophiolite in Chile (Fig. 1), which formed at a trench-proximal Chile Rise (Karsten et al., 1996), is a type example of C-MORB ophiolite. It was emplaced into the South American continental margin as a result of a ridge-trench

collision (Anma et al., 2009). Mid-ocean-ridge ophiolites, in general, correspond to class II and III types in Miyashiro’s (1975) classification of ophiolites based on the presence of tholeiitic and alkaline volcanic rocks. Plume-type (P) ophiolites may form close to plume-proximal spreading ridges and as part of oceanic plateaus (e.g., Caribbean Plateau; Kerr et al., 2009). They have thick plutonic and volcanic sequences (Coffin and Eldholm, 2001; Kerr et al., 2009), and show depleted (D-MORB) to enriched (E-MORB) traceelement patterns (Pearce, 2008). Suprasubduction-zone (SSZ) ophiolites (e.g., Mirdita, Albania; Samail, Oman; Troodos, Cyprus; Fig. 1) form in the extending upper plates of subduction zones, as in the modern Izu-Bonin-Mariana and Tonga-Kermadec arctrench rollback systems (Fig. 1; Hawkins, 2003; Reagan et al., 2010). They may evolve in ex-

tending, embryonic backarc to forearc environments (BA-FA), forearc settings (FA), and both oceanic and continental backarc basins (OBA and CBA, respectively; Table 2). The Rocas Verdes ophiolites in southern Chile are the best examples of suprasubduction-zone continental backarc basin ophiolites (Saunders et al., 1979; Stern and de Wit, 2003). Suprasubduction-zone ophiolites commonly have a Penrose-type structural architecture and may show a MORB–IAT– boninitic geochemical sequence of igneous activity. Suprasubduction-zone forearc ophiolites result from oceanic crust generation during the closure of ocean basins and mark major subduction initiation events (Casey and Dewey, 1984; Dilek and Furnes, 2010; Pearce and Robinson, 2010). The age range among their various ophiolitic subunits is commonly less than 10 m.y. (Dilek and Furnes, 2009). They correspond to the class I ophiolites of Miyashiro

TABLE 1. REPRESENTATIVE EXAMPLES OF MAIN OPHIOLITE TYPES, THEIR GEOGRAPHIC LOCATIONS, APPROXIMATE AGES, AND RELATED REFERENCES ( continued ) Ophiolite Suprasubduction-zone type ( continued ) 9 Magnitogorsk 1

S Urals, Russia

10 11A 11B

Baimak-Buribai Trinity 1 Solund-Stavfjord

SW Urals, Russia California, USA SW Norway

12 13A 13B

Kudi-Kunlun Thetford Mines Bay of Islands

NW China Canada Canada

460–470 479 484

13C 14A

Betts Cove Karmøy

Canada SW Norway

489 474–493

14B 14C

Gulfjellet Leka

SW Norway NW Norway

489 497

15 Volcanic-arc type 1 2A

Lachlan

SE Australia, Tasmania

Itogon Coast Range and Great Valley 1 Zedong 1 Coast Range and Great Valley 2 Zedong 2 Smartville Josephine

Philippines California, USA

30 140

Tibet, China California, USA

127–140 155

Malpas et al. (2003) Shervais et al. (2004), Hopson et al. (2008)

Tibet, China California, USA Oregon and California, USA

155–162 155–165 162–164

5 6A 6B 7A 7B Accretionary type 1 2 3

D’Aguilar 1 D’Aguilar 2 Magnitgorsk 2 Magnitgorsk 3 Trinity 2

E Australia E Australia S Urals, Russia S Urals, Russia California, USA

Malpas et al. (2003) Saleeby et al. (1989), Dilek et al. (1990, 1991) Saleeby et al. (1982), Harper and Wright (1984), Harper et al. (1994) Harper (2003a, 2003b) Spaggiari et al. (2003, 2004) Spaggiari et al. (2003, 2004) Spadea et al. (2003) Spadea et al. (2003) Brouxel et al. (1989), Metcalf et al. (2000)

Mineoka Tokoro Peri-Caribbean 3

4 5 6 7 8 9 10 11

Tamba Solonker 1 Solonker 2 Ganychalan 1 Ganychalan 2 Ganychalan 3 Ganychalan 4 Ganychalan 5

Central Japan Japan Hispaniola, Guatemala, Aruba-Curacao, Central Cuba Japan Central Asia Central Asia NE Russia NE Russia NE Russia NE Russia NE Russia

2B 3A 3B 4A 4B

Location

Age (Ma) 385–400 420 440 440

495–510

360 380 370 385 385 25 60 88–90 135 240 250 420 440 460 480 500

References Spadea and Scarrow (2000), Spadea et al. (2003) Spadea and D’Antonio (2006) Spadea and Scarrow (2000) Brouxel et al. (1989), Metcalf et al. (2000) Furnes et al. (1982); Pedersen (1986), Dunning and Pedersen (1988) Pedersen and Furnes (1991), Furnes et al. (1990, 2003, 2006) Wang et al. (2001, 2002), Yang et al. (1996) Hebert and Laurent (1989), Page et al. (2009), Schroetter et al. (2003) Casey et al. (1985), Suhr (1992), Bedard and Hebert (1996) Varfalvy et al. (1997), Kurth-Velz et al. (2004) Coish et al. (1982), Bedard et al. (1998), Bedard (1999) Furnes et al. (1980), Pedersen (1986), Dunning and Pedersen (1988) Pedersen and Hertogen (1990), Pedersen and Furnes (1991) Furnes et al. (1982), Dunning and Pedersen (1988), Heskestad et al. (1994) Prestvik (1974), Pedersen (1986), Dunning and Pedersen (1988) Pedersen and Furnes (1991), Furnes et al. (1988, 1992) Spaggiari et al. (2003, 2004) Encarnacion (2004) Shervais et al. (2004)

Hirano et al. (2003), Takahashi et al. (2003), Ogawa and Takahashi (2004) Taira et al. (1988), Isozaki (1996) Donnelly (1989), Kerr et al. (1997b), Sinton et al. (1998) Nakae (2000), Koizumi and Ishiwatari (2006) Xiao et al. (2003), Chen et al. (2009) Xiao et al. (2003), Chen et al. (2009) Sokolov et al. (2003) Sokolov et al. (2003) Sokolov et al. (2003) Sokolov et al. (2003) Sokolov et al. (2003v


393


, ) ) 3 9 0 7 0 9 2 1 ( ( . l , ) n a 7 o t 8 h e t l 9 d 1 E r ( a d n d n e o a d G ) n d r , n u i ) e t L 3 r e 0 e S ) d , 0 h ) 0 n 2 a 9 0 ( t 0 r s 7 . i l 2 ( a a e l c 9 1 . c t l ( u e n a . e l i r t A a e t e , ) d ) u f l 5 9 a e u r e 8 0 n s i d r m 9 0 u 2 E n e 1 ( , ) , a d Y ) . , n ) , 2 l 8 6 u 1 0 a 9 8 6 ) t a 9 0 8 e ) 9 9 0 0 1 S ) 9 2 0 0 ( 1 ) 2 , 0 1 ( ( ) 0 i n 9 y ( . l 9 . l . 2 8 ( e 9 h a 9 l 9 . a ( l l 9 7 a t T 9 . 1 t 1 t 9 1 a l h e ( e 1 ( t e a c ( r d ( n e r t s o n d d s l r e n a r n e s e g e y k a r a e k t o t n t z a e p d n u r e h T l a i l i t S a p e u l i v e D R D L B F E E M

S E R ) ) U 8 8 ) 9 0 G ) 9 I 9 0 3 9 F 2 1 ( ( 9 9 1 . . l E 9 ( l 1 H a a . s t t ( l T e e e . a l c t r e i a e D ) u n i t l b N o o n e e 6 n A s 9 a d v p i t T 9 a n n t u 1 X m o a a o ( G E . d m l , ) ) M T R ) e a o ) 0 , , t 6 7 ) t E H 0 7 ) e e 4 3 , 0 9 H 9 9 4 c 8 9 ) r 2 T ) ) ) ( 1 n 9 9 6 8 r 7 1 ( ( e 7 a a . 1 e 1 0 K l l N . . r ( ( 9 I 9 6 7 l 0 , a a . 1 . l e l 1 9 9 a ( 2 f t ( ) D t 9 9 ( 7 a l . t e a t . E 1 n e a e ( 1 ( l R t y e o k e t S e i a 8 e a . l . t 9 s l e v n U s o c a t i a a e 1 l c l ( l g a u d t t n S l e u r e r e e i a e o e n n r N a a t c r r v m M o r c r r s v O t g a e a I t n a l e i e e e a T O V F B S K L K K H K A I V E R b B . N B . l 9 2 9 3 7 9 5 b 2 9 A o a 3 1 2 1 1 4 n - 1 D N a Y h N T A , S . i E . l 6 9 3 7 2 1 0 4 7 9 2 T C o a / 2 3 1 6 1 3 1 8 1 1 0 n 1 R N a V U O S i . t . l A o 1 7 2 1 9 7 1 6 a l T u 1 3 1 3 1 1 1 4 A N n M a D O T . n . l e 7 6 9 3 9 2 1 S o 0 5 7 4 3 6 2 2 3 1 1 1 8 1 8 1 E N a n w 1 1 3 1 5 o C a B N E R a E i e l a F i b e y h E l m y t S C a y l R o l I l . a t , t o a n D a I S a , I C i , N e s s d n e , b b , a A e l n o d e d m i i d i b , u t a r i n l n r i r s o S a u u n s n a l a I l c i N i o C g n s C g o I i n e O e i L L L , I r e p a n a o l l T P n r i a s A a d a c a i a a A u n c e o t a c i n q o u n r n h C r r g t s a a a s e a p c l r r r t r e t O b t e a i o e m o o u y x l a a L n E N C I c M I T G W J C A C R I E H s T n A , o F P D P B B B B B i S t E R P P T R R R R R a M Z B i I S A P O O O B O v C R R R O O P Y O S Z e O M O M M M O M M r T b S M N D E C T M M T b S S A U R C C I t l N a A s a E b C ) ) e O s t s g / t E d n n i r l T e e t I L n a m m e l a s e e e O l I a g g t l e e e c b H i d e d a i r o e e r g l l P s e b b n a i g n i O i t d i d t b t t a d . i r a l l e a a e r 2 e n c e p a p a m p c a o g s s d n y t E o - i a m a t e a d m - e d t L o e s i r c b o d a e i c B c b u o n c c n e t n n e m n r - l A i l o r m g i g i o a c i T d n a z a l i a e e i d e i d m d s i a e c g a r r c h n i h m m o t - t - t m a o r e i n r l i o b i - n n n n n g m l t f x d a a a o x i a a i d t d c t m o i ( e ( e c a i e l o s r r o u n n c a r i c c m p l d o d o l y i o a d c o t d n p l l / e e t b t a e e h a e d h d t s u r a i e n i a t c m e - s k i l i i c i s s l e c m m n n n a c n r p e r o t i u m m o e a i u u a r a - l l r o e n r r m h n e p B d P P T N D E C T l u i p o c u O C M P O S

394

5 7 3 2 5 2

1 3 7 2 6 1 1 4 0 8

y e d a l n w r i h a o l C d N n n n r r e y n u o t e t h e a f s u k r m w u e e o T O N W S

A A A B F B C Z O Z S S Z S S S S

c r a c k r c a a k b c l a a b t c n c i e r n i a n a t e e r c n o o F O C


) 1 9 9 1 ( . l a t e k e l i D

6 0 4

6 3 3

9 6 3 3 9

1 8 5 1

2 2

3 8 2

0 5 2 4 1 2

1 4 7 2 7 6 3 4 0 6 1 8

) b 3 0 0 2 , a 3 0 0 2 , 8 8 9 1 ( . l a t e r e p r a H , ) 4 8 9 1 ( r e p r a H

3 6 3 4 5 9

2 0 9 1

, a , i a n i r n f o r A f o i l i a s , S l e s U a C n e , C , i a p d n n d p a o r a i l i c t e v g t h s i n s e P a h , C s e w A N n h a h a A r o t t r S r S z r W U e U u o o i L N N S

A V

c r a c i n a c l o V

s e s y l a n a f o r e b m u n l a t o T


(1975) and harzburgite-type (HOT) ophiolites of Ishiwatari (1985) and Boudier and Nicolas (1985), which are the products of high degrees of melting of depleted, harzburgitic mantle. Both suprasubduction-zone oceanic backarc basin and continental backarc basin ophiolites form as a result of seafloor spreading in “ensimatic” and “ensialic” settings (respectively). Volcanic-arc (VA) ophiolites form in ensimatic arc settings (e.g., the Philippines, SE Asia; Sierra Nevada, California). They have a polygenetic crustal architecture with a deformed, older oceanic basement, mafic lower crust composed of gabbroic plutons and hypabyssal intrusions, moderately to well-developed dioritic-tonalitic middle crust, andesitic to rhyolitic extrusive rocks and dikes (locally sheeted) forming the upper crust, and volcaniclastic cover (locally subaerial). These crustal units display tholeiitic to calc-alkaline geochemical signatures. Volcanic-arc ophiolites differ from suprasubduction-zone ophiolites based on their thicker and more fully developed arc crust with calc-alkaline compositions. The age range among various ophiolitic subunits in volcanicarc ophiolites can be longer than 20–30 m.y. (Dilek et al., 1991). Accretionary-type ophiolites, occurring in subduction-accretion complexes of active margins, contain fragments of any of the previously outlined ophiolite types and are locally associated with pelagic-hemipelagic sedimentary rocks and trench-fill sediments that may have been deposited on them prior to and after their incorporation into the accretionary prism. These ophiolites may have diverse lithological assemblages, metamorphic grades, styles of deformation, and chemical affinities with no genetic links between them, since they consist of tectonic slices of oceanic rocks scraped off from downgoing plates (e.g., Mineoka ophiolite in central Japan; Ogawa and Takahashi, 2004). They become progressively younger in age structurally downward within subduction-accretion complexes. We do not treat these ophiolites separately in our discussion here because they do not show a distinctive lithological construction, and hence they lack a unique geochemical fingerprint. Geochemical Fingerprinting of Ophiolite Types

We use a selection of diagrams to characterize the geochemical signatures of some wellpreserved examples of the types of ophiolites distinguished here. These diagrams are based on an extensive database (compiled from our own analytical work and the extant literature) that is summarized in Table 2. The literature we used

in our ophiolite classification and geochemicaltectonic fingerprinting is presented in the GSA Data Repository.1 Since lavas and dikes in ophiolites are, in general, subject to various degrees of hydrothermal alteration and greenschist- to amphibolites-facies metamorphism in intra-oceanic conditions, it is important to use elements that are relatively stable during such processes in order for us to determine their primary geochemical compositions. Several studies have been carried out on the element behavior of magmatic rocks that were variably altered and metamorphosed. In general, the mobility of an element relates to the water-rock interactions during reaction (e.g., Bickle and Teagle, 1992). Low-temperature experimental studies of reaction between basalt and seawater have demonstrated minor leaching of Fe and Si and enrichment of Na and Mg; on the other hand, Al, Ti, and P are the least mobile elements, and Ca is variably depleted (Scott and Hajash, 1976; Seyfried et al., 1978). The trace elements Y, Zr, Nb, V, Cr, Co, Ni, rare earth elements (REEs), Th, and Ta are generally relatively immobile (Coish, 1977; Hellman et al., 1979; Shervais, 1982; Seyfried and Mottl, 1982; Dickin and Jones, 1983; Dungan et al., 1983; Mottl, 1983; Staudigel and Hart, 1983; Seyfried et al., 1988; Gillis and Thompson, 1993). A study on the behavior of transition metals (Ti, V, Ni, Cr, Co, Cu, Zn, Fe, Mn) and Mg in metabasic rocks suggests relatively little mobility during medium to high degrees of metamorphism (Nicollet and Andriambololona, 1980). During hydrothermal alteration of basaltic pillow lavas, Ba shows variable alteration trends (Humphris and Thompson, 1978), and Pb becomes moderately to strongly depleted (Teagle and Alt, 2004). Alteration (palagonitization) of the glass rind of pillow lavas results in enrichment of K, Rb, and Cs, particularly the latter two (Hart, 1969; Staudigel and Hart, 1983). Therefore, we paid particular attention in constructing the geochemical diagrams presented here to use those elements that are relatively stable during hydrothermal alteration. In Bowen diagrams (Fig. 3) demonstrating the compositional variability in upper-crustal units (lavas and dikes), the subduction-related suprasubduction-zone and volcanic-arc ophiolites show larger variability in SiO2 and TiO2 at given MgO contents than the subduction-unrelated continental margin, mid-ocean-ridge, and plume ophiolites. The highest variability with respect to these two elements is represented by GSA Data Repository item 2011131, Data source for geochemistry and tectonics of different ophiolite types used in Tables 1 and 2, and for Figures 3–6, is available at http://www.geosociety.org/pubs/ ft2011.htm or by request to [email protected]. 1

the suprasubduction-zone backarc- to forearctype ophiolites, whereas the suprasubductionzone forearc-type ophiolites show invariably low TiO2 (Fig. 3B). The largest spread in MgO is exhibited by the subduction-unrelated plume-type ophiolites (Fig. 3A). In MORBnormalized multi-element diagrams, the continental margin, mid-ocean-ridge, and plume ophiolites display flat patterns between V and Zr, and an increase toward the most incompatible elements (i.e., Ba, Rb, Cs; Fig. 4A). In the same multi-element diagrams, the patterns of the suprasubduction-zone and volcanic-arc ophiolites display much larger variability; they are generally enriched in the most incompatible, nonconservative elements (Cs, Rb, Th) and show generally negative Ta and Nb and positive Pb and Sr anomalies (Fig. 4B). In a Ti-V discrimination diagram (Shervais, 1982), the continental margin, mid-ocean-ridge, and plume ophiolites straddle the field defined by the ratios between 20 and 50, typical of mid-ocean-ridge basalts (Fig. 5A), whereas the suprasubduction-zone and volcanic-arc ophiolites show a wider scatter of Ti/V ratios between <10 and >50 (Fig. 5B). However, the subtypes of both the subduction-related and subductionunrelated ophiolites demonstrate pronounced differences in their Ti-V distributions. For the subduction-unrelated types, the Ti-V data of the lavas and dikes for the plume subtype hardly overlap with those of the continental margin and mid-ocean-ridge trench-proximal subtypes (Fig. 5A). Similarly, for the subduction-related ophiolite types, the mafic lavas and dikes of the suprasubduction-zone forearc subtype exclusively plot in the boninite field and do not overlap with those of the suprasubduction-zone oceanic backarc basin subtype (Fig. 5B). By far, the suprasubduction-zone backarc to forearc subtype shows the largest range in the Ti-V diagram (Fig. 5B). This dispersion of Ti/V ratios is a result of a large geochemical range from boninite and island-arc tholeiite to MORB magmas that occur in subduction-influenced igneous systems (Shervais, 1982; Dilek et al., 2007; Dilek and Furnes, 2009). In the Nb/Yb versus Th/Yb diagram (Pearce, 2008), the lavas and dikes of the continental margin, mid-ocean-ridge, and plume ophiolites plot within the mantle array (Fig. 6A), whereas those of the suprasubduction-zone and volcanicarc ophiolites show a significant shift away from this mantle array, toward the subduction-related Mariana arc field (Fig. 6B). These five elements (Ti, V, Th, Yb, Nb), which we have used in discriminating possible tectonic settings of ophiolitic magma generation, are most immobile during metamorphism and alteration; therefore, they are most reliable as proxies to differentiate


395


Figure 3. Bowen diagrams showA2. Subduction-unrelated A1. Subduction-unrelated ing the relationships between MgO-SiO2 and MgO-TiO2 for Cont. margin Cont. margin subduction-unrelated ophiolites 70 3 Plume Plume (i.e., continental margin, plume, MOR (PP) MOR (PP) and mid-ocean-ridge types) MOR (PD) MOR (PD) (A1 and A2), and subductionMOR (TP) MOR (TP) related ophiolites (i.e., volcanic ) ) arc and suprasubduction-zone % % . . t t [SSZ] types) (B1 and B2). The 60 2 w w ( ( mid-ocean-ridge type (MOR) 2 2 O O i i is subdivided into three sub S T types, i.e., plume-proximal (PP), plume-distal (PD), and trenchproximal (TP). The supra50 1 subduction-zone type (SSZ) is subdivided into four subtypes, i.e., backarc to forearc (BA-FA), forearc (FA), oceanic backarc (OBA), and continental backarc (CBA). Data sources (listed in 40 0 the GSA Data Repository [see 0 5 10 15 20 25 30 0 5 10 15 20 25 30 text footnote 1]): Continental MgO (wt. %) MgO (wt. %) margin type—Ferrara et al. (1976), Beccaluva et al. (1977), B1. Subduction-related B2. Subduction-related Ottonello et al. (1984), Vannucci et al. (1993), Rampone et al. Volc. arc Volc. arc (1998), Montanini et al. (2008). 70 3 SSZ (BA-FA) SSZ (BA-FA) Plume type—Kerr et al. (1996a, SSZ (FA) SSZ (FA) 1996b, 1997), Hastie et al. (2008). SSZ (OBA) SSZ (OBA) Mid-ocean-ridge types, includSSZ (CBA) SSZ (CBA) ing PP subtype—Sigvaldason (1974), Hemond et al. (1993); ) ) % % 2 PD subtype—Kamenetsky et al. . 60 . t t w (2000); TP subtype—Le Moigne w ( ( 2 2 et al. (1996), Guivel et al. O O i i S T (1999). Volcanic-arc type— Yumul et al. (2000), Evans et al. (1991), Metzger et al. (2002), Harper (1984), Harper (2003a, 50 1 2003b), Harper et al. (1988), Dilek et al. (1991). Suprasubduction-zone types, including BA-FA subtype—Dilek et al. (2008), Lippard et al. (1986), 40 0 Einaudi et al. (2003), Godard 0 5 10 15 20 25 30 0 5 10 15 20 25 30 et al. (2003), Auclair and LudMgO (wt. %) MgO (wt. %) den (1987), Rautenschlein et al. (1985), Taylor (1990), Dilek and Thy (1998, 2009), Y. Dilek (personal observation, 1998). FA subtype—Bédard (1999); oceanic backarc basinsubtype—Furnes et al. (2006, and references therein); and continental backarc basin-subtype—Saunders et al. (1979), Stern and Elthon (1979), Stern (1979, 1980), Elthon (1979).

between subduction-related and other magmas (Shervais, 1982; Pearce, 2008), particularly when utilized together with other informative geochemical techniques and field-oriented regional tectonic constraints. Geochemical characterization of different types of ophiolites allows us to distinguish two 396

major groups, one related to or least influenced by subduction-zone processes and the other unrelated to subduction zones. The suprasubduction-zone ophiolites that formed in backarc and incipient arc–forearc tectonic environments (e.g., Mirdita, Albania—Dilek et al., 2007, 2008; Troodos, Cyprus—Robinson et al., 2003;

Pearce and Robinson, 2010), in a forearc setting (e.g., Betts Cove, Canada—Bédard, 1999), and as a volcanic arc (e.g., Smartville, California—Dilek et al., 1991) display the most pronounced variations in geochemical patterns. On the other hand, trench-distal backarc ophiolites that formed in oceanic or continental settings,


Downloaded from gsabulletin.gsapubs.org on February 17, 2011 Ophiolite genesis and global tectonics 100

Cont. margin Plume MOR (PP) MOR (PD) MOR (TP)

A. Subduction-unrelated

10

B R O M / k c o R

1

0.1 Cs Rb Ba Th

U

Ta Nb K

La Ce Pb Pr Sr

P Nd Zr Hf Sm Eu Gd Ti Tb Dy

Y Ho E r Tm Yb Lu

100

V

Sc Co C r Ni

SSZ (BA-FA) SSZ (FA) SSZ (OBA) Volc. arc (MORB-like) Volc. arc (IAT-bon)

B. Subduction-related

10

B R O M / k c o R

1

0.1 Cs Rb Ba Th

U

Ta Nb K

La C e Pb Pr Sr

P

Nd Zr

Hf Sm Eu Gd Ti Tb Dy

Y Ho Er Tm Yb Lu

V

Sc Co C r Ni

Figure 4. Mid-ocean-ridge-basalt (MORB)–normalized multi-element diagrams, showing average values for subduction-unrelated (A) and subduction-related (B) ophiolites. IAT—island-arc tholeiite; bon—boninite. Different types and subtypes of ophiolites are explained in Figure 3. Normalizing values (in ppm) are: Cs (0.007), Rb (0.56), Ba (6.3), Th (0.12), U (0.047), Ta (0.13), Nb (2.33), K (1079), La (2.5), Ce (7.5), Pb (0.3), Pr (1.32), Sr (90), P (314), Nd (7.3), Zr (74), Hf (2.05), Sm (2.63), Eu (1.02), Gd (3.68), Ti (7614), Tb (0.67), Dy (4.55), Y (28), Ho (1.01), Er (2.97), Tm (0.456), Yb (3.05), Lu (0.455), V (300), Sc (40), Co (40), Cr (275), and Ni (100). The elements have been placed in order of their relative incompatibility with spinel-lherzolite mantle (after Pearce and Parkinson, 1993). Data sources (listed in the GSA Data Repository [see text footnote 1]): Continental margin type—Montanini et al. (2008); plume type—Kerr et al. (1996b, 1997), Hastie et al. (2008); mid-ocean-ridge types, including plume-proximal subtype—Hemond et al. (1993); plume-distal subtype—Kamenetsky et al. (2000); trench-proximal subtype—Le Moigne et al. (1996), Guivel et al. (1999); volcanic-arc type—Harper (2003b); suprasubduction-zone types, including BA-FA subtype—Dilek et al. (2008), Dilek and Thy (1998), Y. Dilek (personal observation, 1998); FA subtype—Bédard (1999); and oceanic backarc basin subtype—H. Furnes (personal observation, 1997).

e.g., the Solund-Stavfjord ophiolite in West Norway (Furnes et al., 2006) and the Rocas Verdes ophiolites in the southernmost Andes, Chile (Saunders et al., 1979; Stern and De Wit, 2003), show weaker geochemical evidence of subduction. The groups of ophiolites that are entirely unrelated to subduction processes are the continental margin, mid-ocean-ridge, and plume ophiolites.

PETROGENESIS OF OPHIOLITE TYPES IN DIFFERENT TECTONIC SETTINGS

Figure 7 depicts the petrogenesis of subduction-related and subduction-unrelated types of ophiolites in different tectonic settings. The petrogenesis of a subduction-unrelated continental margin ophiolite involves slow exhumation and

limited partial melting of subcontinental mantle lherzolite (Fig. 8A) and upwelling asthenosphere in response to lithospheric extension and continental rifting (Fig. 7A1; Rampone et al., 2005; Piccardo et al., 2009). Multiple intrusions of MORB-type magma form small olivine gabbro pods and dikes (Fig. 8A) and cause basaltic eruptions on the seafloor (Figs. 7A2 and 8B). Extensional tectonics and associated faulting


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Figure 5. Geochemical data from the subduction-unrelated (A) and subduction-related (B) ophiolite types and their subtypes (see Fig. 3 for explanation) plotted in Ti-V discrimination diagrams (after Shervais, 1982). Data sources (listed in the GSA Data Repository [see text footnote 1]): Continental margin type—Vannucci et al. (1993), Ferrara et al. (1976), Montanini et al. (2008), Beccaluva et al. (1977); plume type—Kerr et al. (1996a, 1996b, 1997), Hastie et al. (2008); mid-ocean-ridge types, including plume-proximal subtype—Sigvaldason (1974), Hemond et al. (1993); plume-distal subtype—Kamenetsky et al. (2000); trench-proximal subtype—Le Moigne et al. (1996), Guivel et al. (1999); volcanicarc type—Evans et al. (1991), Metzger et al. (2002), Harper (1984), Harper (2003a, 2003b), Harper et al. (1988); suprasubduction-zone types, including BA-FA subtype—Dilek et al. (2008), Einaudi et al. (2003), Godard et al. (2003), Dilek and Thy (1998), Y. Dilek (personal observation, 1998); FA subtype—Bédard (1999); oceanic backarc basin-subtype—Furnes et al. (2006, and references therein). Typical Ti/V ratios in C1 and C2 are: 10–20 for island arc; 20–50 for MORB; 20–30 for mixed mid-ocean-ridge basalt (MORB) and island arc, and 1–50 for backarc basins. The boninite field is drawn on the basis of the geochemical data from Crawford (1989).

and shearing may cause tectonic brecciation of the lavas (Fig. 8C). Oceanic crust formation at oceanic spreading axes involves decompression melting of uprising asthenosphere and focused upward ascent of the melt into a melt lens and associated crystal mush 398

zone (Fig. 7A1). Magma injection into a narrow, ~250-m-wide region (Rubin and Sinton, 2007) above the melt lens causes crustal accretion via diking and eruption on the seafloor along the ridge axis. Lavas and dikes have compositions more depleted in incompatible trace elements

than magmas generated from primitive mantle. Locally, melts derived from incompatibleelement–enriched mantle sources may segregate and rise to form off-axis intrusions and to feed near-ridge, E-MORB lavas. Studies of core samples from modern ocean ridges have shown that variations in rates of magma supply and the thermal structure beneath the spreading axis control the mode of magmatic accretion and the architecture of oceanic crust produced. A low and episodic supply of magma to a slow-spreading ridge creates a “cold” environment in which extensional faulting and crustal attenuation accompany seafloor spreading. Amagmatic extension can result in exhumation of serpentinized uppermantle peridotite on the seafloor, and highly thinned lower crust and sheeted dikes (Fig. 7A2; Cannat et al., 1995; Dick et al., 2006). On the other hand, a voluminous supply of magma and the existence of a crustal melt lens at shallow depth (Fig. 7A1) beneath fast-spreading ridges create a “hot”’ environment, in which continuous magma emplacement keeps pace with seafloor spreading. Contemporaneous extension and diking produce Penrose-type oceanic crust underlain by a <1-km-thick transitional Moho (TZ in Fig. 7A2). Intermediate-spreading oceanic crust is similar in structure, but it may have a comparatively thinner volcanic sequence with more pillowed lava flows and a thicker sheeted dike complex (Fig. 7A2; Dilek, 1998). Plume ophiolites form at plume-proximal oceanic ridges or as oceanic plateaus when batches of basaltic and picritic magma originating in a plume head are repeatedly added to preexisting oceanic crust (Fig. 7A1; Coffin and Eldholm, 2001). Lavas range in composition from N-MORB, through T-MORB, to E-MORB. High mantle potential temperatures associated with the plume head result in the highest degree of melting (Pearce, 2008; Hastie and Kerr, 2010). Therefore, we can potentially differentiate plume ophiolites from ophiolites with MORB geochemical affinities by higher Mg contents (Figs. 3A1 and 3A2), resulting from higher degrees of melting, as well as by their internal structure and distinctive volcanic stratigraphy (Fig. 7A2). Typically, a plumerelated ophiolite is characterized by massive basaltic lava flows with subordinate pillowed flows, the occurrence of picritic basalts, and minor sedimentary deposits, all intruded by gabbroic plutons and sills and locally by ultramafic sills (Fig. 7A2; Kerr et al., 1998; Coffin and Eldholm, 2001). Pillow breccias, hyaloclastites, and subordinate sedimentary rocks (chert, shale, and limestone) are locally intercalated with basaltic lava flows at higher stratigraphic levels. The genesis of suprasubduction-zone ophiolites involves the initiation of subduction, followed


Downloaded from gsabulletin.gsapubs.org on February 17, 2011 Ophiolite genesis and global tectonics 10

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Figure 6. Geochemical data from the subduction-unrelated (A) and subduction-related (B) ophiolite types and their subtypes (see Fig. 3 for explanation) plotted in Nb/Yb-Th/Yb discrimination diagram (after Pearce, 2008). OIB—ocean-island basalt; E- and N-MORB—enriched and normal mid-ocean-ridge basalt. Data sources (listed in the GSA Data Repository [see text footnote 1]): Continental margin type—Vannucci et al. (1993), Rampone et al. (1998), Montanini et al. (2008); plume type—Klaver (1987), Kerr et al. (1997), Hastie et al. (2008); midocean-ridge types, including plume-proximal subtype—Hemond et al. (1993); plume-distal subtype—Kamenetsky et al. (2000); trench-proximal subtype—Le Moigne et al. (1996), Guivel et al. (1999); volcanic-arc type: Metzger et al. (2002), Harper (2003a, 2003b); suprasubduction-zone types, including BA-FA subtype—Dilek et al. (2008), Einaudi et al. (2003), Godard et al. (2003), Dilek and Thy (1998), Y. Dilek (personal observation, 1998); FA subtype—Bédard (1999); oceanic backarc basin subtype—H. Furnes (personal observation, 1997)

by rapid slab rollback leading to extension and seafloor spreading in the upper plate (Fig. 7B1). In the subduction initiation stage, magma is first produced by decompressional melting of deep and fertile lherzolitic mantle and produces the earliest crustal units with MORB-like compo-

sitions (Figs. 8D–8F). Fluids derived from the subducted slab have little influence on melt evolution at this early stage. The subsequent phases of melting are strongly influenced by slab dehydration and related mantle metasomatism, melting of subducting sediments, repeated

episodes of partial melting of metasomatized peridotites, and mixing of highly enriched liquids from the lower fertile source with refractory melts in the melt column beneath the extending protoarc-forearc region (Fig. 7B1). Melt aggregation, mixing, and differentiation


399


can take place at many levels within this melt column, and repeated melting of the hydrated mantle leaves behind a highly depleted, olivineand orthopyroxene-rich source. This subarcforearc melting column produces island-arc tholeiite magma that is emplaced into and forms lavas overlying crustal units with MORB-like compositions. Rising temperatures in the mantle wedge, triggered by increased asthenospheric diapirism and lateral flow of hot mantle (the slab edge effect of Pearce and Robinson, 2010) and further influx of slab-derived fluids result in shallow partial melting of the ultrarefractory peridotites (harzburgites), forming Mg- and silica-rich, hydrous, boninitic melts. Replacement of the primary olivine by orthopyroxene (opx) grains in the peridotites and the presence of hydrous minerals (i.e., amphibole), as observed in most of the suprasubduction-zone ophiolites, indicate that the orthopyroxenite forms by the reaction of the preexisting olivine with these boninitic melts (Umino and Kushiro, 1989; Dilek and Morishita, 2009; Morishita et al., 2010). The orthopyroxenite thus represents a reaction product between the migrating melt and the host peridotite in the upper mantle, whereas the harzburgite is the residual, depleted peridotite of the partial melt that produced the orthopyroxenite (Fig. 8G). It is likely, therefore, that geochemical features of boninitic melts are acquired as a result of interaction of migrating melts with depleted peridotites in the mantle wedge (Varfalvy et al., 1997). The harzburgitedunite-orthopyroxenite suite in the upper-mantle peridotites of suprasubduction-zone ophiolites are melting residues and melt migration pathways in the mantle wedge during the incipient stage of arc construction. Boninitic dikes and lavas commonly represent the youngest rock units crosscutting and overlying the earlierformed igneous suites in suprasubduction-zone forearc ophiolites (Figs. 8H and 8K–8L). Suprasubduction-zone ophiolites hence generally display a characteristic, sequential evolution of MORB to island-arc tholeiite to boninitic igneous activity, which manifests itself in a vertically and laterally well-developed chemostratigraphy (Fig. 7B2; Dilek and Furnes, 2009), as also observed in the modern Izu-Bonin-Mariana forearc system (Reagan et al., 2010). The initial stage of construction of a volcanicarc ophiolite involves basic magma. With continued subduction and infiltration of arc magmas, the hydrated mafic crust is partially melted to form tonalitic magmas, and this tonalitic crust grows in thickness as the volcanic arc matures (Fig. 7B1). Residual mafic crust can be transformed into peridotitic restite, and consequently the Moho becomes a fossil melting front (Tatsumi et al., 2008). Volcanic-arc 400

ophiolites thus consist of an older oceanic lithospheric foundation overlain by a mature arc suite, complete with gabbroic plutons and massive diabase in the mafic lower crust, dioritic to tonalitic middle crust, and andesitic to rhyolitic lavas, dike intrusions, and pyroclastic and volcaniclastic rocks in the upper crust (Fig. 7B2). The construction of a volcanic arc is a result of prolonged subduction (~20–40 m.y.) not terminated by colliding continental blocks, as is the case in the evolutionary history of suprasubduction-zone ophiolites (Dilek and Flower, 2003). Sheeted dikes (Figs. 8I–8J) are tabular intrusions of magma flowing laterally and vertically along fractures produced by spreading-related tensile stresses, and they form along a narrow axial zone beneath central rifts along ocean ridges and above subduction zones. The existence of sheeted dikes in ophiolites is conventionally interpreted as strong evidence for the origin of ancient oceanic crust now exposed on land by seafloor spreading (Gass, 1990; Moores and Vine, 1971) and is generally regarded as an essential component of an ophiolite. However, the generation of a sheeted dike complex requires a delicate balance between the rates of spreading and magma supply for a sustained period such that sufficient melt is produced to keep pace with extension in the rift zone (Robinson et al., 2008). In the upper plates of subduction zones, the extension is a consequence of the rate of slab rollback exceeding the rate

of plate convergence, whereas the magma supply is related to the temperature profile and the abundance and nature of fluids in the mantle wedge, the age and lithological makeup of the subducting slab, and the history and extent of melting in the mantle source (Kincaid and Hall, 2003; Robinson et al., 2008). It is rare for the balance between spreading and magma supply rates to be maintained in a suprasubductionzone setting of oceanic crust formation. In the absence of this balance, a sheeted dike complex will not form fully, or even at all, and may instead be replaced by magmatic inflation and the emplacement of plutons, underplating the extrusive sequence (where the rate of magma supply exceeds the spreading rate), or by amagmatic tectonic attenuation of the oceanic crust (where the spreading rate exceeds the rate of magma supply). This phenomenon may explain the scarcity of sheeted dike complexes in nearly 90% of the world ophiolites (Robinson et al., 2008), and should be considered in interpretations of the architecture of putative ancient oceanic crust, particularly in Archean greenstone belts. Continental margin, mid-ocean-ridge, and plume ophiolites may show pronounced variations in trace-element abundances, particularly for the most incompatible elements, which may be related to both different degrees of melting and mantle fertility, but which do not define any particular geochemical evolutionary trend (Fig. 7A3).

Figure 7 ( on following page). Tectonic settings and processes of subduction-unrelated (A1) and subduction-related (B1) ophiolite types, columnar sections depicting simplified structural architecture of the ophiolite types (A2–B2), and generalized changes in element concentration during their evolution (A3–B3). Note that the scale varies from the crust to the mantle in B1. Panels A3 and B3: For subduction-unrelated types (continental margin [CM], mid-ocean-ridge [MOR], and plume [P]), there is no distinct, regular change with time. There may be large (for the most incompatible elements) to moderate (less incompatible to compatible elements) changes in the element concentrations, as indicated by the vertical arrows. For the subduction-related ophiolites, there is a distinct element change from the youngest to the oldest components of the ophiolites. The blank horizontal arrows pointing in opposite directions in B3 indicate that the compositions of mid-ocean-ridge basalt (MORB)–like to island-arc tholeiite (IAT) to boninitic may change to lower or higher contents of the elements indicated. Abbreviations: A1 (CM-type): U. Crust—upper crust; L. Crust—lower crust; Serp. perid.—serpentinized peridotite; A1 (P-type): Cont.—continent; B1: MORB—mid-ocean-ridge basalt; IAT—island-arc tholeiite; BON—boninite. A2 (CM type): Serpt. perd.—serpentinized peridotite; Serp. breccia—serpentinized breccia; P—pillow lava; Lhz—lherzolite; Ol-gabbro—olivine gabbro; A2 (MOR type): Interm.—intermediate; Neovolc.—neovolcanic; TZ—transition zone; M—Moho; DF—detachment fault. A2 (P type): Gb—gabbroic to komatiitic intrusions; ultr. sill—ultramafic sill; picr. bas.—picritic basalt; plw breccia—pillow breccia. B2 (suprasubduction-zone type): MORB, IAT, BON; same as in B1; And.—andesitic lava; Trndj. N—trondhjemite intrusions. B2 (volcanic-arc type): Rhy.—rhyolite; And. lava—andesitic lava; Gran./ton.— granite/tonalite plutons; Gb—gabbro; Di—diorite; DM—depleted mantle; L, M, and HREE—light, middle, and heavy rare earth elements.



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Figure 7.


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Figure 8 ( on this and following page). Field photos from continental margin and various suprasubduction-zone ophiolites, de picting their internal structure and the crosscutting relationships of different ophiolitic subunits. (A) Lherzolitic peridotites of the Jurassic In-Zecca ophiolite (continental margin type) in the Ligurian ophiolites (eastern Corsica) intruded by irregular olivine gabbro dikes and veins. (B) Pillow lavas with normal mid-ocean-ridge basalt (N-MORB) geochemical affinities, resting directly on serpentinized peridotites of the In-Zecca ophiolite. (C) Tectonically brecciated pillow lavas (in B), showing cataclastic shearing in and around the pillow-shaped flows. (D) Layered gabbro rock in the 493 Ma Karmøy ophiolite (suprasubduction-zone backarc to forearc [BA-FA] type) in western Norway intruded by basaltic dikes (D1) with MORB affinities that are in turn crosscut by boninitic dikes (D2). (E) Sheeted dike–gabbro transition zone (Karmøy ophiolite), where leucocratic gabbros and basaltic dikes show mutually intrusive relationships in a Penrose-type crustal pseudostratigraphy. (F) Pillow lavas with island-arc tholeiite (IAT) geochemical affinities in the Karmøy ophiolite crosscut by an island-arc tholeiite dike.

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c i t i a n v i a n l o B

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Suprasubduction-zone and volcanic-arc ophio- incompatible elements in the magmas produced. lites show a characteristic geochemical evo- (2) The mantle melt source becomes enriched in lution. In the early stages of their formation, highly incompatible, nonconservative elements magmas are MORB-like, but during repeated (particularly Cs, Rb, Ba, Th, U) transported in episodes of melting, their mantle source be- subduction-derived fluids and/or felsic melts. comes progressively depleted in the most incompatible elements. The geochemical evolu- Application to Precambrian tion of suprasubduction-zone and volcanic-arc Greenstone Belts ophiolitic magmas is characterized by low abundances of incompatible elements (Cs, Rb, Ba, U, We have selected three Precambrian greenTa Nb, and light [L] REEs) in basaltic andesites, stone belts ranging in age from Paleoproteroandesites, and dacites, which commonly occur zoic (Jormua, Finland) to Neoarchean (Wawa, in the upper parts of their extrusive sequences, Canada) and Paleoarchean (Isua, Greenland), and in young crosscutting dikes in sheeted dike for the purpose of comparing the published geocomplexes. With repeated melting, the residual chemical data for the volcanic and subvolcanic mantle source is progressively enriched in oli- rocks of these sequences with the Phanerozoic vine and orthopyroxene, the principal hosts of ophiolite types as classified herein. compatible elements such as Ni, Co, Cr, and Sc. At a later stage in the magmatic evolution Isua Supracrustal Belt The mafic-ultramafic units of the ca. 3.8 Ga of suprasubduction-zone ophiolites, there is a change from depletion to enrichment in incom- Isua supracrustal belt in Greenland occur in two patible element contents in the younger igneous major tectonostratigraphic units, namely the unrocks relative to MORB; the more incompatible differentiated amphibolites (UA) and Garbenan element is, the more pronounced its enrich- schiefer amphibolites (GA) (e.g., Nutman et al., ment becomes in many suprasubduction-zone 1984, 1997; Rosing et al., 1996; Komiya et al., ophiolite lavas. This phenomenon suggests that 1999; Furnes et al., 2007, 2009). The undifferenthe mantle source undergoes enrichment of tiated amphibolites unit contains all major lithohighly mobile elements during or before the ex- logical units of a typical Penrose-type, complete traction of MORB-like magmas from it. It is the ophiolite sequence, whereas the Garbenschiefer nonconservative, highly incompatible elements, amphibolites unit is composed dominantly of Cs, Rb, Th, and U, that show the most pro- volcaniclastic and volcanic rocks that are comnounced change from depletion to enrichment monly found in immature island arcs. during the late-stage evolution (Fig. 7B3); the other highly incompatible but conservative ele- Wawa Greenstone Belts ments, such as Ta and Nb, remain depleted (e.g., The 2.7 Ga Wawa greenstone belt of the Pearce and Parkinson, 1993). Pb and Sr seem to Superior Province in Canada consists of Albe enriched at an earlier stage than the other non- undepleted and Al-depleted komatiites and conservative incompatible elements, and these Mg- and Fe-tholeiites (Polat et al., 1998, 1999). elements, particularly Pb, increase in concentra- Compositionally, these mafic volcanic and tion from the island-arc tholeiite magmatic stage plutonic rocks are comparable to Phanerozoic to the final boninite activity (Fig. 7B3). ocean plateau basalts that subsequently were Enrichment of the source mantle in slab- tectonically imbricated with primitive arc baderived, nonconservative elements is a com- salts (Polat et al., 1998, 1999). plex process that may involve fluids released from altered oceanic crust and its sedimentary Jormua Complex cover and felsic magmas generated by partial The 1.95 Ga (Peltonen et al., 1996) mafic to melting of subducted sediments (Pearce and ultramafic rocks of the Jormua Complex (JC) Parkinson, 1993; Hawkesworth et al., 1997; occur in the central part of an early ProteroMacdonald et al., 2000; Elburg et al., 2002). zoic (2.3–1.92 Ga) metasedimentary sequence Thus, during the generation of subduction- that is surrounded by Archean basement rocks related ophiolites, two dominant, contempora- in northeastern Finland (Kontinen, 1987; neous processes operate to continuously modify Peltonen et al., 1996). The Jormua Complex the source region and are responsible for the includes pillow lavas and volcanic breccias, typical trace-element patterns of the magmas a sheeted dike complex, mafic cumulates, and produced: (1) Repeated episodes of partial melt- upper-mantle peridotites, and it is tectonically ing progressively deplete the mantle source in disrupted into several blocks. The thickness of incompatible elements and enrich it in compat- the Jormua Complex varies, and in places the ible elements. Inhomogeneities in the mantle lava sequence rests directly upon the uppersource and variable degrees of partial melting mantle rocks, typical of the Ligurian ophiolites could also result in variable concentrations of in the Apennines. The crustal architecture of the 404

Jormua Complex is reminiscent of that seen in slow-spreading oceanic crust and in continental margin ophiolites (Peltonen et al., 1996, 2003). Summary

In the Bowen diagrams (MgO-TiO2), the younger Garbenschiefer amphibolites of the Isua supracrustal belt plot exclusively in the field of subduction-related ophiolites, whereas the undifferentiated amphibolites plot both in the subduction-related and subduction-unrelated fields (Fig. 9A). The Wawa and Jormua metabasalts plot predominantly in the fields of plume and continental margin types of the subductionunrelated ophiolites, respectively (Figs. 9B and 9C). In the multi-element diagrams, the undifferentiated amphibolites of Isua plot within the field of subduction-related ophiolites and display their characteristic features, such as positive Pb anomalies, negative Nb and Ta anomalies, and strong enrichment of Ba and Th. On the other hand, the Garbenschiefer amphibolites show strong depletion of the middle (M) REEs, a typical feature of boninites (Fig. 10A). The Wawa and Jormua metabasalts plot within the field defined by the subduction-unrelated ophiolites and display the same features of flat to moderately enriched patterns as the incompatibility of the elements increase (Figs. 10B and 10C). In the Ti-V discrimination diagram, the Isua data plot in two distinct fields, with the Garbenschiefer amphibolites exclusively in the boninite field (Ti/V < 10), whereas the undifferentiated amphibolites have Ti/V ratios of 20–30 (Fig. 11A) in the mixed MORB and island-arc fields (Shervais, 1982). The volcanic and dike rocks of the Wawa and Jormua sequences, on the other hand, plot entirely within the plume and continental margin types, respectively, of subductionunrelated ophiolites (Figs. 11B and 11C). In the Nb/Yb-Th/Yb discrimination diagram, all the Isua data plot in the subduction-related field (Fig. 12A), whereas the Wawa and Jormua data plot in the subduction-unrelated field (Fig2. 12B and 12C). The Wawa data define a large spread between N-MORB and oceanic-island basalt (though mostly between N-MORB and E-MORB), while the Jormua data cluster tightly around E-MORB (Figs. 12B and 12C). The geochemical character of the metavolcanic and intrusive rocks of the three selected Precambrian greenstone belts indicates that they originated in different tectonic environments. Thus, compared with the geochemical evolution of Phanerozoic ophiolites, the Paleoarchean Isua rocks most likely represent a suprasubduction-zone forearc basin subtype ophiolite, as suggested by Furnes et al. (2009). The Neoarchean Wawa greenstone belt, on the other hand, is more akin to the structural and geo-


Downloaded from gsabulletin.gsapubs.org on February 17, 2011 Ophiolite genesis and global tectonics 3.5

A. Isua, Greenland (3.8 Ga) 3

) 2.5 % t w ( 2 2 O i T 1.5

Garbenschiefer amphibolites Undifferentiated amphibolites Subduction-related Subduction-unrelated

1 0.5

Figure 9. Bowen diagrams showing relationships between MgO and TiO2 for three Precambrian greenstone belts: (A) Isua, Greenland (3.8 Ga), (B) Wawa, Canada (2.7 Ga), and (C) Jormua, Finland (1.95 Ga). The data sources to the enveloping lines for the subduction-related and subduction-unrelated ophiolites, as well as plume (B) and continental margin (C) subtypes, are given in Figure 3. Data sources: Isua, Greenland—Polat et al. (2002), Polat and Hofmann (2003), Komiya et al. (2004), Furnes et al. (2007, 2009); Wawa, Canada—Polat et al. (1999); Jormua, Finland—Kontinen (1987), Peltonen et al. (1996).

0 0

5

10

15

MgO (wt%)

20

25

30

35

3.5

B. Wawa, Canada (2.7 Ga) 3

) 2.5 % t w ( 2

Wawa greenstone Subduction-unrelated

2

O i T1.5

Plume

1 0.5 0 0

5

10

15

MgO (wt%)

20

25

30

35

3.5

C. Jormua, Finland (1.95 Ga) 3

) 2.5 % t w ( 2

Jormua Subduction-related

2

Subduction-unrelated

O i T1.5

Cont. margin

1 0.5 0 0

chemical character of plume-type ophiolites, in agreement with the interpretations of Polat et al. (1999). The early Proterozoic Jormua Complex resembles, both structurally and geochemically, continental margin–type ophiolites, consistent with the interpretations of Peltonen et al. (2003). CONCLUSIONS

Ophiolites are diverse in their internal structure, geochemical makeup, and emplacement mechanisms, and they form in different tec-

5

10

15

MgO (wt%)

tonic environments during the Wilson cycle evolution of ancient ocean basins from rift-drift and seafloor spreading stages to subduction initiation and closure phases. Mafic-ultramafic to felsic rock assemblages that originally formed in different tectonic settings may eventually become nested in collision zones, forming distinct ophiolite complexes with significant diversity in their structural architecture, geochemical fingerprints, and emplacement mechanisms. Differences in the magmatic and structural architecture of ophiolites result from their prox-

20

25

30

35

imity to plumes or trenches, rates and geometry of spreading, mantle temperatures and fertility, and the availability of fluids in the tectonic setting of formation during their primary igneous evolution. Ophiolites are broadly subgrouped into subduction-related and subductionunrelated types. Subduction-related ophiolites include suprasubduction-zone and volcanic-arc types, whereas those unrelated to subduction zones include continental margin, mid-oceanridge (plume-distal and trench-distal), and plume-type ophiolites.


405

Downloaded from gsabulletin.gsapubs.org on February 17, 2011 Dilek and Furnes 1000

Undiff. amph.

A. Isua, Greenland (3.8 Ga)

Garbensch. amph. 100

Subduction-related (max) Subduction-related (min)

B10 R O M / k c o 1 R

0.1

0.01 Cs Rb Ba T h

U

Ta Nb

K

La

Ce Pb

Pr

Sr

P

Nd

Zr

Hf Sm Eu Gd

Ti

Tb Dy

Y

Ho

Er

Tm Yb

Lu

V

Sc Co

Cr

Ni

1000

B. Wawa, Canada (2.7 Ga)

Subduction-unrelated (max) Subduction-unrelated (min)

100

B10 R O M / k c o 1 R

0.1

0.01 Cs Rb Ba T h

U

Ta Nb

K

La

Ce Pb

Pr

Sr

P

Nd

Zr

Hf

Sm Eu Gd

Ti

Tb D y

Y

Ho

Er

Tm Yb

Lu

V

Sc C o

Cr

Ni

1000

C. Jormua, Finland (1.95 Ga)

Subduction-unrelated (max) Subduction-unrelated (min)

100

B10 R O M / k c o 1 R

0.1

0.01 Cs

Rb

Ba

Th

U

Ta

Nb

K

La

Ce

Pb

Pr

Sr

P

Nd

Zr

Hf

Sm

Eu

Gd

Ti

Tb

Dy

Y

Figure 10.

406


Ho

Er

Tm

Yb

Lu

V

Sc

Co

Cr

Ni


Figure 10 ( on previous page). Mid-ocean-ridge basalt (MORB)–normalized multi-element diagrams for the mafic lavas and dikes of the Precambrian greenstone belts in Isua (Greenland), Wawa (Canada), and Jormua (Finland). Normalizing values and the data for drawing the maximum and minimum envelopes for subduction-related and subduction-unrelated ophiolites are provided in Figure 4. Data sources: Isua, Greenland—Polat et al. (2002), Polat and Hofmann (2003), Furnes et al. (2009); Wawa, Canada—Polat et al. (1999); Jormua, Finland—Peltonen et al. (1996).

A. Isua, Greenland (3.8 Ga) 10 20

600

30

400

V

Boninite Subduction-related Subduction-unrelated Garbensch. amph. Undiff. amph.

200

50 0 0

5000

10,000

15,000

20,000

Ti (ppm)


600

400 V

200

Boninite Subduction-related Subduction-unrelated Wawa

0 0

5000

10,000

15,000

20,000

Ti (ppm)


600

400

V

Boninite Subduction-related Subduction-unrelated Cont. margin Jormua

200

0 0

5000

10,000

15,000

20,000

Ti (ppm)

Figure 11. Geochemical data of mafic lav as and dikes from the Precambrian greenstone belts in Isua (Greenland), Wawa (Canada), and Jormua (Finland) plotted in Ti-V discriminant diagrams. Data sources for the enveloped fields for subduction-related and subduction-unrelated ophiolites are given in Figure 5. Data sources: Isua, Greenland—Polat et al. (2002), Polat and Hofmann (2003), Furnes et al. (2007, 2009); Wawa, Canada—Polat et al. (1999); Jormua, Finland—Kontinen (1987), Peltonen et al. (1996).


407

Downloaded from gsabulletin.gsapubs.org on February 17, 2011 Dilek and Furnes 10

Subduction-unrelated

A. Isua, Greenland (3.8 Ga)

Subduction-related Garbensch. amph.

1

OIB

Undiff. amph.

b Y / h T

E-MORB

0.1

N-MORB Figure 12. Geochemical data of mafic lavas and dikes from the Precambrian greenstone belts in Isua (Greenland), Wawa (Canada), and Jormua (Finland) plotted in Nb/Yb-Th/Yb discriminant diagrams. Data sources for the enveloped fields for subduction-related and subduction-unrelated ophiolites, as well as further information for these diagrams, are given in Figure 6. Data sources: Isua, Greenland—Polat et al. (2002), Polat and Hofmann (2003), Furnes et al. (2009); Wawa, Canada— Polat et al. (1999); Jormua, Finland—Peltonen et al. (1996). E- and N-MORB—enriched and normal mid-ocean-ridge basalt; OIB—ocean-island basalt.

0.01 0.01

0.1

1

10

100

Nb/Yb 10


1

b Y / h T 0.1

Subduction-unrelated Subduction-related Wawa greenstone

0.01 0.01

0.1

1 Nb/Yb

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100

10


1

b Y / h T 0.1

Subduction-unrelated Subduction-related Jormua

0.01 0.01

0.1

1

10

100

Nb/Yb

Characterizing ophiolites by their lithological assemblage, internal architecture, and chemical compositions facilitates the identification of the specific tectonic setting of ophiolite generation, which in turn helps us to deduce the processes by which these oceanic rocks were incorporated into continental margins. This new classification of ophiolites provides an effective template for examining the nature of cogenetic relationships between the various parts of ophiolite sequences and for determining the nature of ancient tectonic settings in which the ophiolites formed, particularly for Archean Earth. The application 408

of the ophiolite classification presented here may provide a new conceptual framework to examine potential vestiges of Proterozoic and Archean oceanic lithosphere. We can then use this delineation to better understand the nature of tectonic processes and heat production and dissipation during the Archean. ACKNOWLEDGMENTS

Constructive and thorough comments on earlier versions by Robert Gregory, Brian Robins, and Paul Robinson helped us improve the paper. Our work on ophiolites around the world has been

generously supported by grants from the National Science Foundation, North Atlantic Treaty Organization (NATO) Science Program, Miami University, and the Norwegian Research Council over the years, which we gratefully acknowledge. We wish to thank our colleagues Z. Garfunkel, G. Harper, R. Hébert, E.M. Moores, A. Polat, J. Pearce, R. Pedersen, M. Pubellier, P.T. Robinson, J. Shervais, R. Stern, P. Thy, and J. Wakabayashi for stimulating discussions on various aspects of ophiolites. J. Bédard, B. Murphy, and A. Polat provided objective and insightful reviews of the manuscript, for which we are grateful. We thank Editor Brendan Murphy for inviting us to write this review article for the GSA Bulletin and for his editorial assistance in all stages during the preparation of this paper.


Downloaded from gsabulletin.gsapubs.org on February 17, 2011 Ophiolite genesis and global tectonics REFERENCES CITED

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