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Downloaded from gsabulletin.gsapubs.org on January 27, 2011
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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© 2011 Geological Society of America
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INVITED REVIEW ARTICLE
Ophiolite genesis and global tectonics: Geochemical and
tectonic fingerprinting of ancient oceanic lithosphere
Yildirim Dilek1,† and Harald Furnes2
1
Department of Geology, Shideler Hall, Miami University, Oxford, Ohio 45056, USA, and Faculty of Earth Sciences, China University
of Geosciences at Wuhan, Wuhan 430074, Hubei Province, China
2
Department of Earth Science & Centre for Geobiology, University of Bergen, Bergen 5007, 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: [email protected]
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.
GSA Bulletin; March/April 2011; v. 123; no. 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
387
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Dilek and Furnes
mid-ocean-ridge
Indonesian belt (Cenozoic)
Western Pacific and Cordilleran belts
(Paleozoic-Tertiary)
Alpine - Himalayan belt
(Jurassic - Cretaceous)
°
75
15°W
Appalachian - Caledonian Hercynian - Uralian & Central
Asian belts (early Paleozoic)
Tasmanides (Paleozoic)
165°E
75°E
Sunda Tren
ch
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 marked 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 defini-
388
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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Ophiolite genesis and global tectonics
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
Geological Society of America Bulletin, March/April 2011
389
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Dilek and Furnes
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
Geological Society of America Bulletin, March/April 2011
Number of major ophiolites
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Ophiolite genesis and global tectonics
Central Asian ophiolites
Age (Ma)
Ng
Pg
Tertiary
Cretaceous
Jurassic
Triassic
Permian
Mesozoic
Carb.
Devonian
Sil.
Ord.
Camb.
Paleozoic
Major events to which
ophiolites are related
Age (Ma)
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.
Geological Society of America Bulletin, March/April 2011
391
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Dilek and Furnes
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
Location
Age (Ma)
References
Continental margin type
1
Tihama
2
Ligurian
Red Sea, Saudi Arabia
Italy
20
200
310
320
410
3
4
5
Mid-ocean-ridge type
1A
1B
Ust-Belaya 1
Ust-Belaya 2
Nurali
NE Russia
NE Russia
S Urals, Russia
Macquarie Isl.
Taitao
SW Pacific
S Chile
2
Khoy
Iran
4
Masirah
5
Horo Kanai
6
Kuyul 1
7
Kuyul 2
8
Kuyul 3
9
Nurali
Plume type
1A
Loma de Hiero
1B
Bolivar
2
Nicoya
3
Peri-Caribbean 1
4
Peri-Caribbean 2
5
Duarte
6
Loma La Monja
7
Mino-Tamba 1
8
Mino-Tamba 2
Suprasubduction-zone type
1
Zambales
2
Antique
3A
Troodos
3B
Semail
W Indian Ocean
Central Hokkaido, Japan
NE Russia
NE Russia
NE Russia
S Urals
3C
10
10
98-103
150
165–180
190
200
210
405
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)
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)
Philippines
Panay, Philippines
Cyprus
Oman
40–44
75–80
92–94
92–95
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)
(continued)
392
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Ophiolite genesis and global tectonics
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
Location
Age (Ma)
References
Suprasubduction-zone type (continued)
9
Magnitogorsk 1
S Urals, Russia
385–400
Spadea and Scarrow (2000), Spadea et al. (2003)
Spadea and D’Antonio (2006)
10
Baimak-Buribai
SW Urals, Russia
420
Spadea and Scarrow (2000)
11A
Trinity 1
California, USA
440
Brouxel et al. (1989), Metcalf et al. (2000)
11B
Solund-Stavfjord
SW Norway
440
Furnes et al. (1982); Pedersen (1986), Dunning and Pedersen (1988)
Pedersen and Furnes (1991), Furnes et al. (1990, 2003, 2006)
12
Kudi-Kunlun
NW China
460–470
Wang et al. (2001, 2002), Yang et al. (1996)
13A
Thetford Mines
Canada
479
Hebert and Laurent (1989), Page et al. (2009), Schroetter et al. (2003)
13B
Bay of Islands
Canada
484
Casey et al. (1985), Suhr (1992), Bedard and Hebert (1996)
Varfalvy et al. (1997), Kurth-Velz et al. (2004)
13C
Betts Cove
Canada
489
Coish et al. (1982), Bedard et al. (1998), Bedard (1999)
14A
Karmøy
SW Norway
474–493
Furnes et al. (1980), Pedersen (1986), Dunning and Pedersen (1988)
Pedersen and Hertogen (1990), Pedersen and Furnes (1991)
14B
Gulfjellet
SW Norway
489
Furnes et al. (1982), Dunning and Pedersen (1988), Heskestad et al. (1994)
14C
Leka
NW Norway
497
Prestvik (1974), Pedersen (1986), Dunning and Pedersen (1988)
Pedersen and Furnes (1991), Furnes et al. (1988, 1992)
15
Lachlan
SE Australia, Tasmania
495–510
Spaggiari et al. (2003, 2004)
Volcanic-arc type
1
Itogon
Philippines
30
Encarnacion (2004)
2A
Coast Range and
California, USA
140
Shervais et al. (2004)
Great Valley 1
2B
Zedong 1
Tibet, China
127–140
Malpas et al. (2003)
3A
Coast Range and
California, USA
155
Shervais et al. (2004), Hopson et al. (2008)
Great Valley 2
3B
Zedong 2
Tibet, China
155–162
Malpas et al. (2003)
4A
Smartville
California, USA
155–165
Saleeby et al. (1989), Dilek et al. (1990, 1991)
4B
Josephine
Oregon and California, USA
162–164
Saleeby et al. (1982), Harper and Wright (1984), Harper et al. (1994)
Harper (2003a, 2003b)
5
D’Aguilar 1
E Australia
360
Spaggiari et al. (2003, 2004)
6A
D’Aguilar 2
E Australia
380
Spaggiari et al. (2003, 2004)
6B
Magnitgorsk 2
S Urals, Russia
370
Spadea et al. (2003)
7A
Magnitgorsk 3
S Urals, Russia
385
Spadea et al. (2003)
7B
Trinity 2
California, USA
385
Brouxel et al. (1989), Metcalf et al. (2000)
Accretionary type
1
Mineoka
Central Japan
25
Hirano et al. (2003), Takahashi et al. (2003), Ogawa and Takahashi (2004)
2
Tokoro
Japan
60
Taira et al. (1988), Isozaki (1996)
3
Peri-Caribbean 3
Hispaniola, Guatemala,
Aruba-Curacao, Central Cuba 88–90
Donnelly (1989), Kerr et al. (1997b), Sinton et al. (1998)
4
Tamba
Japan
135
Nakae (2000), Koizumi and Ishiwatari (2006)
5
Solonker 1
Central Asia
240
Xiao et al. (2003), Chen et al. (2009)
6
Solonker 2
Central Asia
250
Xiao et al. (2003), Chen et al. (2009)
7
Ganychalan 1
NE Russia
420
Sokolov et al. (2003)
8
Ganychalan 2
NE Russia
440
Sokolov et al. (2003)
9
Ganychalan 3
NE Russia
460
Sokolov et al. (2003)
10
Ganychalan 4
NE Russia
480
Sokolov et al. (2003)
11
Ganychalan 5
NE Russia
500
Sokolov et al. (2003v
Geological Society of America Bulletin, March/April 2011
393
TABLE 2. OPHIOLITE/OCEANIC CRUST TYPES, THEIR LOCATIONS, AND REFERENCES TO DATA SOURCES, AND ABBREVIATIONS USED IN THE TEXT AND THE FIGURES
No.
No.
No.
No.
anal.
anal.
anal.
anal.
Ophiolite/oceanic crust type
Abbreviations
Location
Bowen
Multi
V/Ti
Th-Yb-Nb
Reference to data sources
Continental margin
CM
Internal Ligurides,Italy
27
2
Ottonello et al. (1984), Rampone et al. (1998)
External Ligurides, Italy
26
11
26
19
Vannucci et al. (1993), Montanini et al. (2008)
North Apennine, Italy
39
39
Ferrara et al. (1976)
Corsica
13
13
Beccaluva et al. (1977)
Mid-ocean ridge
MOR
Plume-proximal mid-ocean ridge
MOR PP
Iceland
119
37
67
39
Sigvaldason (1974), Hemond et al. (1993)
Plume-distal mid-ocean ridge
MOR PD
Macquarie Island
12
12
12
12
Kamentsky et al. (2000)
Trench-proximal mid-ocean ridge
MOR TP
Taitao Peninsula, S. Chile
31
31
31
9
Le Moigne et al. (1996), Guivel et al. (1999)
Normal mid-ocean ridge basalt
NMORB
Depleted (in the incompatible elements)
DMORB
mid-ocean-ridge basalt
Enriched (in the incompatible elements)
EMORB
mid-ocean-ridge basalt
Crustally contaminated mid-ocean-ridge basalt
CMORB
Transitional mid-ocean-ridge basalt
TMORB
Plume
P
Gorgona Island, Colombia
10
10
Kerr et al. (1996a)
Western Colombia
85
19
84
23
Kerr et al. (1997a)
Jamaica
17
17
17
17
Hastie et al. (2008)
Curacao, Caribbean Sea
84
11
19
19
Klaver (1987), Kerr et al. (1996b)
Ocean-island basalt
OIB
Suprasubduction zone
SSZ
Backarc to forearc
SSZ BA-FA
Albania
113
46
102
45
Dilek et al. (2008)
Cyprus
56
Rautenschlein et al. (1985), Auclair and Ludden (1987),
Taylor (1990)
Turkey
61
40
61
25
Dilek and Thy (1998, 2009)
Oman
134
15
113
57
Lippard et al. (1986), Einaudi et al. (2003), Godard et al. (2003)
Forearc
SSZ FA
Newfoundland
47
22
47
23
Bedard (1999)
Oceanic backarc
SSZ OBA
Western Norway
802
802
Furnes et al. (2006, and references therein)
Continental backarc
SSZ CBA
Southern Chile
67
Elthon (1979), Saunders et al. (1979), Stern and Elthon (1979),
Stern (1980)
Volcanic arc
VA
Luzon, Philippines
53
39
Evans et al. (1991), Yumul et al. (2000)
North Cascades,
6
6
6
Metzger et al. (2002)
Washington, USA
93
22
93
40
Harper (1984), Harper et al. (1988, 2003a, 2003b)
Northwestern California,
USA
Sierra Nevada, California,
4
Dilek et al. (1991)
USA
Total number of analyses
1902
283
1581
336
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Ophiolite genesis and global tectonics
(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
1
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]
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
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Dilek and Furnes
TiO2(wt. %)
SiO2 (wt. %)
TiO2 (wt. %)
SiO2 (wt. %)
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., volcanicarc and suprasubduction-zone
[SSZ] types) (B1 and B2). The
60
2
mid-ocean-ridge type (MOR)
is subdivided into three subtypes, 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),
B2. Subduction-related
B1. 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);
PD subtype—Kamenetsky et al.
60
2
(2000); TP subtype—Le Moigne
et al. (1996), Guivel et al.
(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,
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Ophiolite genesis and global tectonics
100
Cont. margin
Plume
MOR (PP)
MOR (PD)
MOR (TP)
A. Subduction-unrelated
Rock/MORB
10
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 Er Tm Yb Lu
V
Sc Co Cr Ni
100
SSZ (BA-FA)
SSZ (FA)
SSZ (OBA)
Volc. arc (MORB-like)
Volc. arc (IAT-bon)
Rock/MORB
B. Subduction-related
10
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 Er Tm Yb Lu
V
Sc Co Cr 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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Dilek and Furnes
A. Subduction-unrelated
600
20
10
30
V
400
50
Cont. marg.
Plume
MOR (PP)
MOR (PD)
MOR (TP)
Boninite
200
0
0
5000
10,000
15,000
20,000
Ti (ppm)
B. Subduction-related
600
V
400
SSZ (BA-FA)
SSZ (FA)
SSZ (OBA)
Volc. arc
Boninite
200
0
0
5000
10,000
15,000
20,000
Ti (ppm)
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
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Ophiolite genesis and global tectonics
10
A. Subduction-unrelated
Cont. crust
Mariana
arc-basin
Th/Yb
1
Cont. marg.
Plume
MOR (PP)
MOR (PD)
MOR (TP)
0.1
0.01
0.01
0.1
1
10
100
Nb/Yb
10
B. Subduction-related
Cont. crust
OIB
1
Th/Yb
Mariana
arc-basin
E-MORB
0.1
SSZ (BA-FA)
SSZ (FA)
SSZ (OBA)
N-MORB
Volc. arc
0.01
0.01
0.1
1
10
100
Nb/Yb
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
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Dilek and Furnes
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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Ophiolite genesis and global tectonics
A1
B1
Subduction - unrelated ocean crust
Subduction - related ocean crust
Suprasubduction - Zone (SSZ) type
Depth (km)
Postrift
Synrift sediments
0
0
10
10
20
20
Subcontinental
30
100
50
40
km
100
50
0
120 km
30
Asthenosphere
40
20 km
Continental margin (CM) type
Mid-ocean ridge (MOR) type
Shallow
intrusion
nonconservative
arc (VA) type
Depth
V olcanic
Backarc
10
km
Plume (P) type
120 km
fluid flow
300°C
Partial
melt.
zone
600°C
900°C
1200°C
0
0
100
100
1000
Depth
(km)
1000
2000
2000
4000
4000
A2
B2
P type
MOR type
CM type
Serpt.
perd.
Serp. breccia/
ophicalcite
Chert
Fast
Interm.
SL
P
DF
dikes
Gb
0.5
km
A3
young
M
pluton
TZ
massive
lava
Gb
ultr. sill Undifferentiated
Ocean Crust
0.5 km
andesite
sheeted
dikes
Depleted mantle
Depleted mantle
0.5 km
Gb
0.3 km
15 km
Plume source
SL
MORB-like
And.
lava
Gran./ ton.
Depleted mantle
Basalt
lava
Gb
Di
Gb
Trndj
1 km
Rhy.
Bon.
P
B3
CM - MOR - P types
IAT
dacite
volcaniclastic/
pyroclastic rocks
DM
10 km
Strongly
depleted mantle
DM
SSZ - VA types
10 m.y.
Ol-gabbro
Subcontinental
mantle (Lhz)
picr. bas.
plw breccia
SL
Time
SL
Slow
Neovolc.
zone
VA type
SSZ type
Sea level (SL)
Figure 7.
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Lherzolite
Olv-gabbro
dikes
A
B
D1
Pillow lava
D2
D2
D1
Layered
gabbro
C
D
Gabbro
Dikes
Dike
E
F
Figure 8 (on this and following page). Field photos from continental margin and various suprasubduction-zone ophiolites, depicting 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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G
H
K
D3
Boninitic
sill
D1
D2
I
L
Boninitic
lava
Figure 8 (continued). (G) Clinopyroxene porphyroclast-bearing harzburgite in the Middle Jurassic (165 Ma) Eastern Mirdita ophiolite of
Albania (suprasubduction-zone BA-FA type), crosscut by networks of orthopyroxenite dikes, dikelets, and veins. These intrusions represent boninitic melt channels that migrated upward into the refractory harzburgite (see Dilek and Morishita, 2009; Morishita et al., 2010).
(H) Plastically deformed layered gabbros in the 92 Ma Kizildag ophiolite in southern Turkey (suprasubduction-zone FA type), intruded by
a boninitic sill and a dikelet. (I–J) Sheeted dike swarms (moderately to vertically dipping) in the Kizildag ophiolite. (K) Basaltic andesite
dikes (D1) with an island-arc tholeiite affinity, intruded by plagiogranite dikes (D2), which are in turn crosscut by a late-stage boninitic dike
(D3). (L) Boninitic lavas (“sakalavites”) in the Kizildag ophiolite. See Dilek and Thy (2009) for details.
Cpx-Harzburgite
Orthopyroxenite
dikes & veins
Mylonitic gabbro
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Dilek and Furnes
Suprasubduction-zone and volcanic-arc ophiolites show a characteristic geochemical evolution. In the early stages of their formation,
magmas are MORB-like, but during repeated
episodes of melting, their mantle source becomes progressively depleted in the most
incompatible elements. The geochemical evolution of suprasubduction-zone and volcanic-arc
ophiolitic magmas is characterized by low abundances of incompatible elements (Cs, Rb, Ba, U,
Ta Nb, and light [L] REEs) in basaltic andesites,
andesites, and dacites, which commonly occur
in the upper parts of their extrusive sequences,
and in young crosscutting dikes in sheeted dike
complexes. With repeated melting, the residual
mantle source is progressively enriched in olivine and orthopyroxene, the principal hosts of
compatible elements such as Ni, Co, Cr, and Sc.
At a later stage in the magmatic evolution
of suprasubduction-zone ophiolites, there is a
change from depletion to enrichment in incompatible element contents in the younger igneous
rocks relative to MORB; the more incompatible
an element is, the more pronounced its enrichment becomes in many suprasubduction-zone
ophiolite lavas. This phenomenon suggests that
the mantle source undergoes enrichment of
highly mobile elements during or before the extraction of MORB-like magmas from it. It is the
nonconservative, highly incompatible elements,
Cs, Rb, Th, and U, that show the most pronounced change from depletion to enrichment
during the late-stage evolution (Fig. 7B3); the
other highly incompatible but conservative elements, such as Ta and Nb, remain depleted (e.g.,
Pearce and Parkinson, 1993). Pb and Sr seem to
be enriched at an earlier stage than the other nonconservative incompatible elements, and these
elements, particularly Pb, increase in concentration from the island-arc tholeiite magmatic stage
to the final boninite activity (Fig. 7B3).
Enrichment of the source mantle in slabderived, nonconservative elements is a complex process that may involve fluids released
from altered oceanic crust and its sedimentary
cover and felsic magmas generated by partial
melting of subducted sediments (Pearce and
Parkinson, 1993; Hawkesworth et al., 1997;
Macdonald et al., 2000; Elburg et al., 2002).
Thus, during the generation of subductionrelated ophiolites, two dominant, contemporaneous processes operate to continuously modify
the source region and are responsible for the
typical trace-element patterns of the magmas
produced: (1) Repeated episodes of partial melting progressively deplete the mantle source in
incompatible elements and enrich it in compatible elements. Inhomogeneities in the mantle
source and variable degrees of partial melting
could also result in variable concentrations of
404
incompatible elements in the magmas produced.
(2) The mantle melt source becomes enriched in
highly incompatible, nonconservative elements
(particularly Cs, Rb, Ba, Th, U) transported in
subduction-derived fluids and/or felsic melts.
Application to Precambrian
Greenstone Belts
We have selected three Precambrian greenstone belts ranging in age from Paleoproterozoic (Jormua, Finland) to Neoarchean (Wawa,
Canada) and Paleoarchean (Isua, Greenland),
for the purpose of comparing the published geochemical data for the volcanic and subvolcanic
rocks of these sequences with the Phanerozoic
ophiolite types as classified herein.
Isua Supracrustal Belt
The mafic-ultramafic units of the ca. 3.8 Ga
Isua supracrustal belt in Greenland occur in two
major tectonostratigraphic units, namely the undifferentiated amphibolites (UA) and Garbenschiefer amphibolites (GA) (e.g., Nutman et al.,
1984, 1997; Rosing et al., 1996; Komiya et al.,
1999; Furnes et al., 2007, 2009). The undifferentiated amphibolites unit contains all major lithological units of a typical Penrose-type, complete
ophiolite sequence, whereas the Garbenschiefer
amphibolites unit is composed dominantly of
volcaniclastic and volcanic rocks that are commonly found in immature island arcs.
Wawa Greenstone Belts
The 2.7 Ga Wawa greenstone belt of the
Superior Province in Canada consists of Alundepleted and Al-depleted komatiites and
Mg- and Fe-tholeiites (Polat et al., 1998, 1999).
Compositionally, these mafic volcanic and
plutonic rocks are comparable to Phanerozoic
ocean plateau basalts that subsequently were
tectonically imbricated with primitive arc basalts (Polat et al., 1998, 1999).
Jormua Complex
The 1.95 Ga (Peltonen et al., 1996) mafic to
ultramafic rocks of the Jormua Complex (JC)
occur in the central part of an early Proterozoic (2.3–1.92 Ga) metasedimentary sequence
that is surrounded by Archean basement rocks
in northeastern Finland (Kontinen, 1987;
Peltonen et al., 1996). The Jormua Complex
includes pillow lavas and volcanic breccias,
a sheeted dike complex, mafic cumulates, and
upper-mantle peridotites, and it is tectonically
disrupted into several blocks. The thickness of
the Jormua Complex varies, and in places the
lava sequence rests directly upon the uppermantle rocks, typical of the Ligurian ophiolites
in the Apennines. The crustal architecture of the
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-
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Ophiolite genesis and global tectonics
3.5
A. Isua, Greenland (3.8 Ga)
TiO2 (wt%)
3
2.5
Garbenschiefer amphibolites
2
Undifferentiated amphibolites
Subduction-related
1.5
Subduction-unrelated
1
0.5
0
0
5
10
15
MgO (wt%)
20
25
30
35
3.5
B. Wawa, Canada (2.7 Ga)
TiO2 (wt%)
3
2.5
Wawa greenstone
Subduction-unrelated
2
Plume
1.5
1
0.5
0
0
5
10
15
MgO (wt%)
20
25
30
35
3.5
C. Jormua, Finland (1.95 Ga)
3
TiO2 (wt%)
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).
Jormua
2.5
Subduction-related
2
Subduction-unrelated
1.5
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.
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Dilek and Furnes
1000
Undiff. amph.
A. Isua, Greenland (3.8 Ga)
Garbensch. amph.
100
Subduction-related (max)
Rock/MORB
Subduction-related (min)
10
1
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
Ho
Er
Tm Yb
Lu
V
Sc
Co
Cr
Ni
1000
B. Wawa, Canada (2.7 Ga)
Subduction-unrelated (max)
Subduction-unrelated (min)
Rock/MORB
100
10
1
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
Ho
Er
Tm Yb
Lu
V
Sc
Co
Cr
Ni
1000
C. Jormua, Finland (1.95 Ga)
Subduction-unrelated (max)
Subduction-unrelated (min)
Rock/MORB
100
10
1
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
Geological Society of America Bulletin, March/April 2011
Ho
Er
Tm
Yb
Lu
V
Sc
Co
Cr
Ni
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Ophiolite genesis and global tectonics
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).
600
10
A. Isua, Greenland (3.8 Ga)
20
30
V
400
Boninite
Subduction-related
Subduction-unrelated
Garbensch. amph.
Undiff. amph.
200
50
0
0
5000
10,000
15,000
20,000
Ti (ppm)
B. Wawa, Canada (2.7 Ga)
600
V
400
200
Boninite
Subduction-related
Subduction-unrelated
Wawa
0
0
5000
10,000
15,000
20,000
Ti (ppm)
C. Jormua, Finland (1.95 Ga)
600
V
400
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 lavas 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).
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Dilek and Furnes
10
Subduction-unrelated
A. Isua, Greenland (3.8 Ga)
Subduction-related
Garbensch. amph.
1
OIB
Th/Yb
Undiff. amph.
E-MORB
0.1
N-MORB
0.01
0.01
0.1
1
10
100
Nb/Yb
10
B. Wawa, Canada (2.7 Ga)
1
Th/Yb
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.
0.1
Subduction-unrelated
Subduction-related
Wawa greenstone
0.01
0.01
0.1
1
Nb/Yb
10
100
10
C. Jormua, Finland (1.95 Ga)
Th/Yb
1
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.
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Ophiolite genesis and global tectonics
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SCIENCE EDITOR: B. MURPHY
ASSOCIATE EDITOR: J. BÉDARD
MANUSCRIPT RECEIVED 15 NOVEMBER 2010
REVISED MANUSCRIPT RECEIVED 14 DECEMBER 2010
MANUSCRIPT ACCEPTED 15 DECEMBER 2010
Printed in the USA
Geological Society of America Bulletin, March/April 2011
411