to understand subduction initiation, study forearc crust: to understand forearc crust, study...

LITHOSPHERE | Volume 4 | Number 6 | www.gsapubs.org 469

RESEARCH

INTRODUCTION

A better understanding of the mechanisms by which new subduction zones form is criti-cal for advancing the solid Earth sciences. Until we can reconstruct how and why this happens, we cannot pretend to understand a wide range of important Earth processes and properties, including lithospheric strength, composition, and density, and the driving force behind plate motions. In spite of this, our understanding of the subduction initiation process has advanced slowly, for two important reasons: (1) Sub-duction initiation is an ephemeral process, so

there are few active examples, and (2) nearly all of the evidence for tectonic, magmatic, and sedimentary responses to subduction initia-tion is preserved in forearcs, which are deeply submerged and buried beneath sediments. We would prefer to study subduction initiation in progress, but there are few places to do this. One such active region however, is the Puysegur subduction zone off the coast of southern New Zealand (LeBrun et al., 2003; Sutherland et al., 2006). However, as only a narrow segment of the Australia-Pacifi c transform plate margin is affected, studies of Puysegur cannot capture all of the processes that accompany major subduc-tion initiation events, i.e., those that change the lithospheric force balance suffi ciently to cause changes in plate motion and stimulate volumi-nous magmatism, as discussed herein. Such episodes shaped the western margin of North America in Mesozoic time (Dickinson, 2004), established a convergent margin along SW Eurasia in Late Cretaceous time (Moghadem and Stern, 2011), and engendered most of the active subduction zones of the western Pacifi c

in Eocene time (Ishizuka et al., 2011). Major subduction initiation episodes are hemispheric in scale and necessarily reorganize upper-man-tle fl ow, and in many cases are accompanied by widespread and voluminous igneous activity.

Here, we outline a strategy that promises to accelerate our understanding of processes asso-ciated with major subduction initiation episodes by considering both the subduction initiation record preserved in forearcs and insights from studying well-preserved ophiolites. The record of subduction initiation is preserved in igneous crust and upper-mantle residues and the associ-ated sediments on the overriding plate next to the trench. These collectively comprise the forearc (Dickinson and Sealey, 1979) and provide the best record of subduction initiation. Signifi cant parts of forearcs may be lost by tectonic erosion (Scholl and von Huene, 2009); nevertheless, whatever remains contains the best record of the processes that accompanied subduction ini-tiation of that particular convergent margin. We explore why this record has been overlooked and summarize recent studies of forearc crust and

To understand subduction initiation, study forearc crust: To understand forearc crust, study ophiolites

R.J. Stern1*, M. Reagan2*, O. Ishizuka3*, Y. Ohara4*, and S. Whattam5*1GEOSCIENCES DEPARTMENT, UNIVERSITY OF TEXAS AT DALLAS, RICHARDSON, TEXAS 75083-0688, USA2DEPARTMENT OF GEOSCIENCE, UNIVERSITY OF IOWA, IOWA CITY, IOWA 52242, USA3INSTITUTE OF GEOSCIENCE AND GEOINFORMATION, GEOLOGICAL SURVEY OF JAPAN/AIST, CENTRAL 7, 1-1-1, HIGASHI, TSUKUBA, IBARAKI 305-8567, JAPAN4HYDROGRAPHIC AND OCEANOGRAPHIC DEPARTMENT OF JAPAN, TOKYO 104-0045, JAPAN5DEPARTMENT OF EARTH AND ENVIRONMENTAL SCIENCES, KOREA UNIVERSITY, SEOUL 136-701, REPUBLIC OF KOREA

ABSTRACT

Articulating a comprehensive plate-tectonic theory requires understanding how new subduction zones form (subduction initiation). Because subduction initiation is a tectonomagmatic singularity with few active examples, reconstructing subduction initiation is challenging. The lithosphere of many intra-oceanic forearcs preserves a high-fi delity magmatic and stratigraphic record of subduction initiation. We have heretofore been remarkably ignorant of this record, because the “naked forearcs” that expose subduction initiation crustal sections are dis-tant from continents and lie in the deep trenches, and it is diffi cult and expensive to study and sample this record via dredging, diving, and drilling. Studies of the Izu-Bonin-Mariana convergent margin indicate that subduction initiation there was accompanied by seafl oor spread-ing in what ultimately became the forearc of the new convergent margin. Izu-Bonin-Mariana subduction initiation encompassed ~7 m.y. for the complete transition from initial seafl oor spreading and eruption of voluminous mid-ocean-ridge basalts (forearc basalts) to normal arc volcanism, perhaps consistent with how long it might take for slowly subsiding lithosphere to sink ~100 km deep and for mantle motions to evolve from upwelling beneath the infant arc to downwelling beneath the magmatic front. Many ophiolites have chemical features that indicate formation above a convergent plate margin, and most of those formed in forearcs, where they were well positioned to be tectoni-cally emplaced on land when buoyant crust jammed the associated subduction zone. We propose a strategy to better understand forearcs and thus subduction initiation by studying ophiolites, which preserve the magmatic stratigraphy, as seen in the Izu-Bonin-Mariana forearc; we call these “subduction initiation rule” ophiolites. This understanding opens the door for on-land geologists to contribute fundamentally to understanding subduction initiation.

LITHOSPHERE; v. 4; no. 6; p. 469–483 | Published online 16 May 2012 doi: 10.1130/L183.1

*E-mails: [email protected]; [email protected]; [email protected]; [email protected]; [email protected].

Editor’s note: This article is part of a special issue ti-tled “Initiation and Termination of Subduction: Rock Re-cord, Geodynamic Models, Modern Plate Boundaries,” edited by John Shervais and John Wakabaya shi. The full issue can be found at http://lithosphere.gsapubs.org/content/4/6.toc.

For permission to copy, contact [email protected] | © 2012 Geological Society of America

STERN ET AL.

470 www.gsapubs.org | Volume 4 | Number 6 | LITHOSPHERE

upper mantle, and what the results reveal about subduction initiation. The expense and diffi -culty of directly studying forearc igneous rock exposures are huge obstacles to our progress, so we explore the potential of some ophiolites for illuminating forearc composition and magmatic stratigraphy. Ophiolites are exposed on land and so are vastly easier and cheaper to study than forearcs. We conclude that those ophiolites that formed in a forearc provide important opportu-nities for advancing our understanding of sub-duction initiation. The strategy of comparative study of igneous forearc crust and ophiolites, coupled with geodynamic modeling, promises to lead to major advancements in our under-standing of subduction initiation processes.

FOREARCS

Forearcs comprise the bulk of any arc-trench system, occupying the ~150–200-km-wide region above the subducted plate between the trench and the magmatic arc. Forearcs are rela-tively stable and low standing—intra-oceanic forearcs lie entirely below sea level—and are morphologically unimpressive compared to spectacular volcanoes of the fl anking magmatic arc and the tremendous gash of the trench. For these reasons, it is understandable that forearcs were either overlooked or misunderstood when the geologic implications of plate tectonics were fi rst being explored in the late 1960s and 1970s. During this time, thinking about forearcs was dominated by examples on or near continents, such as California, Japan, Alaska, and Indone-sia (for an account of early thinking about what would come to be called forearcs, see Dickin-son, 2001). Even today, the textbook example

of a convergent plate margin is provided by the Late Mesozoic of California, with the Francis-can mélange representing exhumed subduction-zone material, the Great Valley Group repre-senting the forearc basin, and the Sierra Nevada Batholith representing the roots of the magmatic arc. This is indeed an excellent example of a sediment-rich convergent margin, but empha-sis on California and other sediment-dominated forearc examples has inhibited appreciation of forearc crust itself.

Many—but not all—continental forearcs are excellent examples of convergent margins affected by high sediment fl ux. Some continental convergent margins—such as Peru-Chile and NE Japan—do not have high sediment fl ux, but these have not been textbook examples because the interesting outcrops are in very deep water, and thus are diffi cult and expensive to study. In con-trast, forearcs away from continents are mostly sediment starved (Clift and Vannucchi, 2004). Such naked forearcs expose crust and upper mantle, which are readily accessed by drilling through thin sediment cover, as was done during Deep Sea Drilling Project (DSDP) Leg 60 and Ocean Drilling Program (ODP) Legs 125 and 126 in the Izu-Bonin-Mariana arc and ODP Leg 135 in the Tonga forearc (Bloomer et al., 1995).

EROSIVE VERSUS ACCRETIONARY

FOREARCS

Subordinate proportions of forearcs are accretionary, growing by deposition of large sediment loads from a fl anking continent, which bury forearc crust beneath forearc basins and then overfl ow to the trench, where these sedi-ments briefl y ride on the subducting plate before

some is scraped off to form an accretionary prism. Such situations of forearc thickening and widening are globally unusual, because most forearcs lose upper-plate crust to the subduc-tion zone due to tectonic erosion, as a result of normal faulting, oversteepening, and basal frac-turing and abrasion along the plate interface. Another misconception (due to bias toward studying sediment-rich convergent margins) is that all inner-trench slopes have very low slopes (<3°), when, in fact, erosive margins, especially those exposing igneous basement, are much steeper, typically with slopes of 3°–7° (Clift and Vannucchi, 2004). Estimates of the proportion of accretionary versus erosive convergent margins vary. According to Clift and Vannucchi (2004), 57% of the cumulative length of trenches is ero-sive and 43% is accretionary, whereas Scholl and von Huene (2007) estimated that 74% and 26% are erosive and accretionary, respectively. Thickness of sediment on the downgoing plate is the single most important control on whether a margin is erosive or accretionary. A sediment thickness of ~500 m divides the two types of margins. Other factors favoring tectonic erosion include collision of large bathymetric features such as seamounts (Clift and Vannucchi, 2004) and presence of rasping grabens on the downgo-ing plate (Hilde, 1983).

Although most forearcs are erosive, they are more poorly known than accretionary forearcs because they are harder to study, and a smaller research community has been interested in them. Erosive forearcs are exclusively submarine, so studying them requires research vessels with technology to examine and sample the bottom (Fig. 1A). Compared to accretionary forearcs, erosive forearcs lie in deeper water, farther from

Figure 1. Photographs of on-land (left) and submerged (right) forearc exposures. Left photo shows how easy it is for geoscientists to examine litholo-

gies and structures on land. Right photo is taken from Shinkai 6500 YK1012, Dive 1231, ~6000 m deep in the southern Mariana Trench. Only one or two

scientists at a time can go down to examine and sample rocks. Costs of an on-land fi eld trip are a miniscule fraction of the expense of a submarine fi eld

trip. Photo on the left is Franciscan radiolarian chert exposed on the Marin headlands, California (photo by S. Graham). Field of view on right is ~7 m.


Subduction initiation, forearcs, and ophiolites | RESEARCH

the continents, making them more diffi cult and expensive to study. In contrast, accretionary forearcs lie in shallower water, and parts rise above sea level (e.g., Kodiak Island, Alaska; Shimanto Belt, SW Japan; and Nias and Men-tawai, Sumatra), where they are relatively easy to study (Fig. 1B). Accretionary margins are found near continents and so are usually in some nation’s territorial waters or exclusive economic zone, which attracts study by the scientists of that nation. Most of the great accretionary mar-gins lie in the Northern Hemisphere, facing the largest continents, biggest rivers, and greatest sediment fl ux. The Northern Hemisphere is also where the richest nations are—the ones most likely to support large-scale geoscientifi c efforts needed for marine tectonic studies. In contrast, erosive margins often lie far away from conti-nents, rich nations, and scientists. Accretionary

margins contain potential economic deposits of hydrocarbons, which attract the attention of oil companies; erosive margins do not. Finally, accretionary margins attract more geophysical interest than do erosive margins. Accretionary margins are characterized by progressive defor-mation of sedimentary layering, which spatially and temporally changes from fl at lying and unconsolidated at the trench to steeply dipping and lithifi ed arcward (Fig. 2A). Such lithostruc-tural variations reward seismic-refl ection profi l-ing with spectacular images, on which structural interpretations, publications, and proposals can be based. In contrast, intra-oceanic forearcs are generally sediment starved because they are far from continents, and thus large sediment fl uxes. These comprise a subclass of erosive margins known as “naked forearcs” (Stern, 2002), which lack thick sedimentary cover and obvious imbri-

cation. Seismic-refl ection profi ling over naked forearcs yields much less interesting seismic refl ection images, and they are more diffi cult to interpret (Fig. 2B). It is so much easier and more rewarding to study accretionary margins that it is a wonder that erosive margins are studied as much as they are.

In spite of the challenges of studying erosive forearcs, it is important that we do so. Naked forearcs expose igneous infrastructure in the inner trench wall and bury this crust under thin sediments of the forearc itself, providing the best opportunities for in situ investigations of this lithosphere by drilling, dredging, and diving.

FOREARC CRUST AND UPPER MANTLE

An important characteristic of many naked forearcs is the presence of peridotite—exposed

5 km V.E. ~2x

3

10

De

pth

(k

m)

Naked forearc (Mariana)2

7

21

Sediment-floodedtrench

--

V.E. ~1.5x

A

B

Sedimented Forearc (Indonesia)

brabrabranchnchnchnchch thththththrrusrusrusttthtt ttttt

décdécdédé olllllloll meeemem ntntntntnt

back thrust

out-of-sequence thrust(backstop thrust)

reduced coherentimaging

branch thrust

décollement

Figure 2. Comparison of seismic-refl ection profi les of (A) sedimented and (B) naked (unsedimented) forearcs. (A) Prestack depth-migrated section of

multichannel refl ection profi le off the Sunda Strait, from Kopp and Kukowski (2003) and interpreted by them. An arcward increase in material strength

results in a segmentation of the margin. Faint seaward- and landward-dipping faults cut the trench fi ll in the protothrust zone of segment I, indicat-

ing the fi rst stages of faulting. The deformation front marks the onset of faulting in conjugate pairs of fore-thrusts and back thrusts. The frontal active

accretionary prism (segment II) is composed of tilted thrust slices separated by regularly spaced thrust faults. The transition to the fossil accretionary

prism of the outer high is marked by a prominent out-of-sequence thrust. Segment III forms the backstop to the frontally accreted material and dis-

plays much reduced tectonic activity, mainly manifested in the occasional reactivation of previous thrusts, which helps adjust the taper. (B) Mariana

forearc, trench, and part of incoming Pacifi c plate shown as depth section (MCS Line 22–23, from Oakley et al., 2008). Circles locate the points along

the plate refl ection where depths were recorded. Normal faults on the incoming plate and the Mariana forearc are interpreted. The toe of the forearc is

uplifted as the fl ank of a Pacifi c plate seamount subducts. M—seafl oor multiple; V.E.—vertical exaggeration.

STERN ET AL.


upper mantle—in the lower trench wall. Trench peridotite exposures demonstrate that the Moho is also exposed, along with a complete crustal section at shallower depths, and indicate the thickness of the crust.

Peridotite is exposed in the inner walls of intra-oceanic trenches at depths >8 km in the Tonga Trench (Bloomer and Fisher, 1987) but can be found as shallow as 5800 m in the south-ern Mariana Trench (Michibayashi et al., 2009). Such depths are mostly beyond the reach of manned submersibles, which currently cannot descend below 6500 m, so forearc peridotites are rarely sampled except by dredging. Still, we know about intra-oceanic peridotite exposures in four trenches: Izu-Bonin, Mariana, Tonga, and South Sandwich (Figs. 3 and 4). In addi-tion, mantle peridotite is brought up by serpen-tine mud volcanoes, which are common in the Mariana forearc and are also known from the Izu forearc (Fryer, 2002).

Because of its importance for understanding the nature and origin of intra-oceanic forearcs, some basic concepts about mantle peridotite in general and forearc peridotite in particular are presented here. These are residues after partial melting and complement magmatic rocks such as lavas—especially basalts—which are more common subjects of marine petrologic study.

Because Earth has had mantle since shortly after it formed, but this has been modifi ed by melt extraction and mixing with subducted materials, idealized compositions are useful for this discussion. For example, “primitive” mantle (PM) refers to an idealized chemical composi-tion after the core segregated but before the con-tinental crust was extracted. PM is also known

as “bulk silicate earth” (BSE). Several studies have estimated PM compositions, including Mg# (100Mg/[Mg + Fe] = 89–90), CaO (2.8–3.7 wt%), and Al

2O

3 (3.5–4.5 wt%; see table 2 of

Lyubetskaya and Korenaga, 2007). These esti-mates constrain minimum Mg# and maximum CaO and Al

2O

3 contents of the upper-mantle

source region of most basalts. Because Earth has been recycling surface materials and melting to make basalt for several billion years, signifi cant tracts of primitive upper mantle are unlikely to exist. Instead, the concept of “fertile mid-ocean-ridge basalt (MORB)–type mantle” (FMM; Pearce et al., 2000) is more useful. FMM is also

an idealized upper-mantle composition, but one which acknowledges the extraction of the conti-nental crust. “Pyrolite” is another idealized com-position (Green and Falloon, 1998) that is very similar to FMM. FMM and pyrolite approximate the composition of upper mantle that partially melts to generate oceanic crust beneath diver-gent plate boundaries (spreading ridges) and to produce arc melts beneath convergent margins (note that other components such as pyroxenite exist in the mantle, but these melt almost com-pletely, leaving no identifi able residue).

If PM, BSE, FMM, and pyrolite were rocks instead of ideas, they would be classifi ed as

Tonga

Kermadec

NewHebrides

New BritainVanuatu

Mariana

Izu-Bonin

AleutiansKuriles

South Sandwich

LesserAntilles

AegeanCascades

And

es

Indonesia

Makran

Himalaya

Japan

Philippines

Central America

0° 60°E 120°E 180° 120°W 60°W

60°N

30°

30°

0°

60°S

Peridotite

Gabbro

Volcanics

De

pth

be

low

se

a le

vel (

km

)

MORB

Figure 4. Interpretive section of a typical intra-oceanic forearc as reconstructed from dredging of the

Tonga Trench. Figure is from Bloomer and Fisher (1987), reproduced with permission of Journal of Geol-

ogy. N- and E- MORB refer to normal and enriched mid-ocean ridge basalt. V.E.—vertical exaggeration.

Figure 3. Known locations of peri-

dotite exposures in the lower parts

of inner-trench walls, shown in

bold: (1) South Sandwich (Pearce

et al., 2000); (2) Tonga (Fisher and

Engel, 1969; Bloomer and Fisher,

1987); (3) Mariana (Bloomer, 1983;

Ohara and Ishii, 1998; Michibayashi

et al., 2007, 2009); (4) Izu-Bonin

(Okamura et al., 2006; Ishizuka et

al., 2011). Mariana and Izu-Bonin

forearcs also bring up peridotite

in serpentine mud volcanoes

(Fryer et al., 1995; Parkinson and

Pearce, 1998). Peridotites repre-

sent the bases of in situ ophiolites

that make up the crust and upper

mantle of intra-oceanic forearcs.

Other inner-trench walls may also

expose peridotite, but most have

not been sampled as extensively

as these four.



lherzolite, an olivine-rich ultramafi c rock that contains >5% clinopyroxene (cpx), the remain-der being orthopyroxene and an aluminous phase (spinel or garnet; Fig. 5A). Changes in peridotite composition due to partial melting are simple and clear. Melt depletion dimin-ishes abundances of cpx, CaO, and Al

2O

3 and

increases Cr# (Cr/Cr + Al) in the residual spi-nel (Fig. 5). This is because basalts, which are rich in CaO and Al

2O

3 (~12 and 16 wt%,

respectively), are generated by partial melting of lherzolite, which is much poorer in CaO and Al

2O

3 (~3 and ~4 wt%, respectively for FMM;

Fig. 5B). Clinopyroxene contains nearly all the CaO in peridotite, whereas spinel in FMM contains ~35% Al

2O

3 (a few percent Al

2O

3 is

also dissolved in pyroxene). Melt depletion decreases the proportion of cpx, so that the residue progressively changes from lherzolite to harzburgite (<5% cpx); extreme melt depletion yields dunite (>90% olivine; Fig. 5A).

Even though most forearc peridotite expo-sures are serpentinized, there are robust mineral and whole-rock compositional characteristics that are remarkably unaffected by such altera-tion. These include changes in the proportions of minerals (Fig. 5A), major-element bulk chemistry (Fig. 5B), and spinel compositions (Fig. 5C). All of these refl ect the amount of melt depletion, as discussed already.

These approaches allow us to compare the “refractoriness” of peridotites from various tec-tonic settings, i.e., how much melt depletion they have experienced. This reveals that forearc peridotites are the most depleted ultramafi c rocks from any modern tectonic environment, with the highest Cr# spinels and lowest propor-tion of cpx and whole-rock abundances of CaO and Al

2O

3 (Bonatti and Michael, 1989). Not all

forearc peridotites are so depleted; for example, some from the South Sandwich forearc include lherzolites with up to 3.7% Al

2O

3 and 4.4%

CaO, along with spinels with Cr# as low as ~0.4 (Pearce et al., 2000). Morishita et al. (2011) documented two populations of spinel in Izu-Bonin forearc dunites, one group with moder-ate Cr# (0.4–0.6), and the other with high Cr# (>0.8). Variations in spinel compositions not-withstanding, forearc peridotites are dominated by ultradepleted compositions rarely found in other tectonic environments.

The unusually depleted nature of forearc peridotites requires unusual melting conditions: abnormally high temperature, volatile fl ux, or both. Whatever the cause, these depletions are all the more noteworthy because forearcs asso-ciated with mature arcs have unusually low heat fl ow (Stein, 2003) and rarely are volcanically active. Whatever conditions caused the unusu-ally extensive melting beneath forearcs no lon-

ger exist. Such transitory conditions are linked to subduction initiation in the next section.

Igneous rocks of exposed forearc crust above peridotites (Fig. 4) are only now becoming the focus of geoscrutiny. At one time, forearc crust was thought to comprise oceanic crust that was trapped when subduction began (e.g., Dickin-son and Sealey, 1979), so the origin of forearc igneous crust has not, until recently, been much studied. This is changing, partly because of what is now recognized as the unusually strong deple-tion of forearc peridotites and because of inter-est in boninites. Boninites are lavas with unusual combinations of high silica and magnesium coupled with low calcium and aluminum abun-dances. These compositional features refl ect low-pressure melting of harzburgitic mantle (e.g., Falloon and Danyushevsky, 2000), which is not otherwise observed on modern Earth. In addition, boninites are enriched in large ion lithophile elements (LILEs; elements with low valence and large ionic radius) and light rare earth elements (LREEs) relative to high fi eld

strength elements (HFSEs; elements with high valence and small ionic radius). High LILE/HFSE ratios in boninites refl ect metasomatism of the source mantle by hydrous fl uid released from subducted crust and sediments (Gill, 1981; Stern et al., 1991; Pearce et al., 1992). This fl uid lowers peridotite melting temperature at the same time that it re-enriches it in fl uid-mobile elements, including LILEs and LREEs.

Not all naked forearcs have boninite, but at least one of them—the Izu-Bonin-Mariana arc system—does (Fig. 6A; Stern et al., 1991; Macpherson and Hall, 2001). Probably the best exposure of boninite in the world is found in the Bonin (Ogasawara) islands (Taylor et al., 1994). These boninites erupted on the seafl oor ca. 46–48 Ma, shortly after subduction began ca. 52 Ma (Ishizuka et al., 2006, 2011). Recent studies of Izu-Bonin-Mariana forearc crust exposed in the inner-trench wall (Fig. 6B) reveal that boninite may be the uppermost component of forearc crust, underlain by thicker, slightly older tholeiitic basalts, which Reagan et al.

Backarc

basin

peridotite

Mariana forearc peridotite

Abyssal peridotite

Boninite

Forearc peridotite

Abyssal peridotite

Cr#

Mg#

5

15

25

50

60

70

80

9010

20

30

40

50

Ol

Opx Cpx

Dunite

Harzburgite

Lherzolite

FMM

Mel

t dep

letio

n

A B

C

FMM

FMM

MD

Figure 5. What happens to peridotites when

they melt: (A) Changes in modal mineralogy

(and thus rock type) with progressive melting,

plotted on International Union of Geological

Sciences (IUGS) classifi cation for peridotites

(LeBas and Streckeisen, 1991). Fertile mid-

ocean-ridge basalt (MORB)–type mantle (FMM)

is mostly peridotite, consisting of olivine (Ol),

orthopyroxene (Opx), and clinopyroxene (Cpx),

comprising lherzolite. Melting to produce basal-

tic melts depletes peridotite in clinopyroxene,

so that residual peridotite after ~20% melting

is harzburgite, with <5% clinopyroxene. Melting

of harzburgite yields boninite melts and resid-

ual dunite. (B) Bulk-rock abundances of Al2O

3

versus CaO (volatile free, normalized to 100%

total), showing how melting depletes perido-

tites in these elements. The compositions of

FMM and mantle residue after 5%, 15%, and

25% partial melting of FMM are from Pearce

and Parkinson (1993). Fields for abyssal and forearc peridotites are from Pearce et al. (1992). (C)

Composition of spinels in peridotite, plotted on Cr# (Cr/Cr + Al) versus Mg# (Mg/Mg + Fe+2) dia-

gram. Fields are after Dick and Bullen (1984), modifi ed to show the composition of backarc basin

peridotites from the Mariana Trough (Ohara et al., 2002). Melting preferentially extracts Al from

spinel, increasing Cr# as melting progresses. Composition of FMM spinels and approximate trend

of melt depletion (MD) are also shown.

STERN ET AL.


(2010) called “forearc basalts.” Forearc basalt lavas and related dikes have chemical compo-sitions similar to MORB. Forearc basalts were fi rst recognized in the southern Mariana forearc SE of Guam, where they crop out trenchward of thin boninites and younger arc rocks (Rea-gan et al., 2010). This outcrop pattern, as well as the volcanic stratigraphy drilled in the Mari-ana forearc at DSDP Site 458 (Fig. 6B) indi-cates that forearc basalts here are older than the boninites and were likely the fi rst lavas to erupt when the Izu-Bonin-Mariana subduction zone formed (Reagan et al., 2010; Ishizuka et al., 2011). Below the forearc basalts, there are gabbroic rocks, and then mantle peridotites, as summarized in Figure 6B (Ishizuka et al., 2011).

Forearc basalts have major-element compo-sitions that are broadly MORB-like, although signifi cant differences exist, at least for Izu-Bonin-Mariana forearc basalts. Forearc basalts are generally not as rich in TiO

2 as is typical

MORB, which often contains >1.4 wt% TiO2.

Forearc basalts can also be more rich in SiO2

than typical MORB (most forearc basalts con-tain <51% but forearc basalt–boninite tran-sitional lavas at DSDP Sites 458 and 459 are

richer in silica). In spite of these differences, most forearc basalts also have trace-element (Figs. 7B and 7C) and isotopic compositions (Reagan et al., 2010) that are MORB-like, whereas younger forearc lavas have composi-tions suggesting that subducted fl uids were involved in their genesis. For example, lavas generated by melting in the presence of a fl uid from a subducting plate (e.g., Mariana arc lavas and boninites) typically have elevated Th/Yb compared to MORB. On a plot of Th/Yb versus Nb/Yb, most Izu-Bonin-Mariana forearc basalts plot with MORB along the unmodifi ed “mantle array,” whereas younger, subduction-infl uenced lavas trend toward Th/Yb typical of Mariana arc lavas (Fig. 7E). Note that Figure 7 also plots the chemostratigraphies of ophiolitic lavas, a point which is discussed further below.

DeBari et al. (1999) studied the Izu-Bonin-Mariana inner-trench wall (6100–6500 m deep) near 32°N (Fig. 6A) and documented the pres-ence of MORB-like basalts. They interpreted these to represent older oceanic crust that was trapped when subduction began. However, the composition of these lavas is identical to those of forearc basalts from elsewhere in the Izu-Bonin-

Mariana forearc, and we prefer to interpret this as another exposure of basalts that formed when Izu-Bonin-Mariana subduction began.

To conclude this section, it is clear that our understanding of the composition of forearc crust is incomplete, largely due to the diffi culty of accessing and studying this material. Much of what we think we know is based upon studies of the Izu-Bonin-Mariana forearc. We anticipate that Izu-Bonin-Mariana forearc crust is repre-sentative of igneous rocks that form when sub-duction begins, but we cannot be sure until we have studied the igneous rocks of other forearcs in similar detail.

SUBDUCTION INITIATION

In this section, we discuss igneous activity associated with the formation of a new sub-duction zone. As presented already, much of what we understand about the igneous crust of forearcs comes from studies of the Izu-Bonin-Mariana system. The fundamental question we address for this forearc, and one that is pertinent for many others is: Does subduction initiation generate this broad swath of crust, which ulti-mately forms the forearc (Fig. 8)? Also, what happens when the subduction zone itself forms? For the Izu-Bonin-Mariana arc system, we have documented the progression of igneous activity in the forearc about the time that the Pacifi c plate changed its motion, and we have concluded that this activity resulted from the dynamic response of the crust and upper mantle to subduction initi-ation. The basic idea is: at ca. 52 Ma, old, dense lithosphere of the Pacifi c plate began to sink, perhaps due to differential subsidence across a lithospheric weakness, such as an old fracture zone (Fig. 9A; Stern and Bloomer, 1992). We conclude that subduction initiation at this time was hemispheric in scale: much of the western Pacifi c, extending south from Izu-Bonin-Mari-ana to Fiji and the Tonga-Kermadec convergent margin, formed new subduction zones about this time. This was accompanied by voluminous gen-eration of forearc basalts, boninite, and related igneous rocks, much of which is now preserved in these forearcs. We are not sure whether these new subduction zones were caused by, or were the cause of the change in the Pacifi c plate abso-lute motion, from NNW to WNW at ca. 50 Ma, as refl ected by the bend in the Emperor-Hawaii seamount chain (Sharp and Clague, 2006), but both events happened about the same time.

There are many challenges to this summary of Izu-Bonin-Mariana subduction initiation. These include hypotheses that: (1) interaction with a mantle plume was responsible for boninite for-mation (Macpherson and Hall, 2001); (2) extru-sion of Indian Ocean–Asian asthenosphere due

Peridotite

Gabbro/Mesozoicbasalt

Lithology Age(Ma)

50–52

50–52Basalt (FAB)

Boninite(and relateddifferentiates)

44–48

44–45High-Mgandesite

Arc tholeiites and calc-alkalinerocks

37–44

Sheeted dike

Pillow lavahyaloclastite

Approximatedepth (m)

6780

Gabbro

?

?

6300

5500–4760

Subaeriallyexposed

Basalt

Japa

n

Tokyo ∗

140° 150°

20° N

30°

*

.....

.

.

.

....

.. ......

.

**

*

***

500 km

PhilippineSea plate

Pacificplate

Izu-Bonin-M

ariana

A

130° E

arc system

B1

2

3

*

B

Figure 6. (A) Schematic map of the Izu-Bonin-Mariana arc system, showing principal tectonic fea-

tures and forearc crustal sections discussed in text. 1—study of Reagan et al. (2010); 2—study of

Ishizuka et al. (2011); 3—study of DeBari et al. (1999). Asterisk marks location of Deep Sea Drill-

ing Project (DSDP) Site 458; B is location of Bonin (Ogasawara) Islands. White region is seafl oor

<2500 m deep. Dashed line shows location of profi le in Figure 8. (B) Schematic columnar section of

crust exposed in the Bonin inner-trench wall, from Ishizuka et al. (2011), used with permission by

Elsevier. FAB—forearc basalt.



-

A

C

E

B

D

F

Figure 7. Whole-rock geochemical data of stratigraphically constrained (i.e., lower versus upper) subalkaline lavas and latest-

stage dikes of Izu-Bonin-Mariana forearc lavas (left) and ophiolites of the Eastern Mediterranean–Persian Gulf region (right).

Forearc chemical data are from Reagan et al. (2010) and Ishizuka et al. (2011). Ophiolite data are from Albania (Mirdita; data

from Dilek et al., 2008); Greece (Pindos; data from Saccani and Photiades, 2004); Cyprus (Troodos; data from Flower and Levine,

1987); and Oman (Semail; data from Godard et al., 2003) as compiled by Whattam and Stern (2011). (A) Izu-Bonin-Mariana

(IBM) forearc and (B) Tethyan ophiolites on the SiO2 versus FeOt/MgO subalkaline affi nity discrimination diagram (Miyashiro,

1974). (C) Izu-Bonin-Mariana forearc and (D) Tethyan ophiolite lavas on chondrite-normalized rare earth element (REE) plots

(REE concentrations from Nakamura, 1974). (E) Izu-Bonin-Mariana forearc and (F) Tethyan ophiolite lavas on the Nb/Yb ver-

sus Th/Yb plot (Pearce, 2008). Compositional data have been fi ltered to include samples with reported major-oxide totals of

98%–102% and loss on ignition (LOI) <7%. In A, total iron is expressed as FeOt (= Fe2O

3 × 0.89), and oxide concentrations were

recalculated and normalized to 100% on an anhydrous (volatile-free) basis. In E and F, abbreviations are N- and E-M—normal

and enriched mid-ocean-ridge basalt (MORB). Note that in B, the only “lower lava” samples that plot as calc-alkaline (n = 4) are

from Troodos. In addition, Troodos is not represented on F because the data set of Flower and Levine (1987) does not include

concentrations of Th, Nb, and Yb. Modern Mariana arc lava data (blue dots) in E are from Peate and Pearce (1998). FAB—forearc

basalt; FC—fractional crystallization.

STERN ET AL.


to Tethys closure was a major cause of western Pacifi c arc rollback and basin opening (Flower et al., 2001); (3) boninite formation refl ected the intersection of an active spreading ridge and a subduction-transform fault transition (Des-champs and Lallemand, 2003); (4) compression across a preexisting fracture zone was required (Hall et al., 2003); (5) lateral compositional buoyancy contrast within oceanic lithosphere controls subduction initiation (Niu et al., 2003); (6) the new Izu-Bonin-Mariana convergent mar-gin cut across, rather than followed, preexist-ing lithospheric fabric, such as remnant arcs, fracture zones, and spreading ridges (Taylor and Goodliffe, 2004); and (7) subduction of the Pacifi c-Izanagi spreading ridge triggered a chain reaction of tectonic plate reorganizations that led to Izu-Bonin-Mariana and Tonga-Kermadec subduction initiation (Whittaker et al., 2007).

There is clearly a lot of uncertainty about what caused the Izu-Bonin-Mariana subduc-tion initiation and attendant boninite volca-nism, but there is no disagreement that it was accompanied by voluminous igneous activity. This formed Izu-Bonin-Mariana forearc crust (Figs. 6 and 8) as well as crustal tracts well to the west of the present volcanic front, includ-ing the West Mariana Ridge and Kyushu-Palau Ridge. Stern and Bloomer (1992) conserva-tively (i.e., assuming generation of 6-km-thick crust) estimated that 1200–1800 km3 of crust were produced per kilometer of arc during Izu-Bonin-Mariana subduction initiation, and that this episode lasted 10 m.y., for a crustal growth rate of 120–180 km3/km. This is equivalent to the volume of crust produced at a mid-ocean-ridge spreading at 2–3 cm/yr. In fact, Izu-Bonin-Mariana mean crustal thickness produced dur-ing this episode was probably greater than 6 km, as shown in Figure 8 for the Izu forearc. For the

7.0

TrenchVolcanic front

050100150180

25

20

15

10

5

0

V.E. ~4xMantle Vp ~8.0

Subduction

interface

Lower oceanic crust Vp ~7.0

Upper oceanic crust Vp ~6.8 Moho

Subducted Moho

Mantle

Vp ~8.0

7.3

6.46.8 6.5

Gabbro 6.7

6.3 6.35.8

4.8 4.3

4.5 4.5

Altered basalt and diabase

Vp ~1.9 sediments

3.3

3.2Volcanics

Inner forearc Outer forearc

De

pth

(k

m)

Distance from trench (km)

v v v v v v v v v v v

LMLM

Early protoforearc spreading (FAB)

AM

A

B

Oceanic crust

AM

Localized magmatic arc(volcanic front)D

Late protoforearc spreading (VAB/BON)Cv

v v v v v v v v v v

Upwelling fertile asthenosphere; no interaction with slab-derived fluid

Depleted mantle stagnates;strong interaction with slab-derived fluid

Sinking slab; rapid trench rollback

True subduction; trench rollback slows

Forearc crustFig. 11

Young Old

TF/FZ

Figure 9. Subduction initiation, formation of the forearc, and evolution of magmatic systems, modi-

fi ed after Metcalf and Shervais (2008). (A) Older, thicker, colder, and denser lithosphere (right) is jux-

taposed with young, thinner, hotter, and more buoyant lithosphere across a zone of weakness (e.g.,

transform fault or fracture zone [TF/FZ]). (B) Subsidence of old lithosphere allows asthenosphere to

fl ood over it. Upwelling asthenosphere melts due to decompression, generating mid-ocean-ridge

basalt (MORB)–like basalt (forearc basalts of Reagan et al., 2010) accompanied by seafl oor spread-

ing. (C) Continued lithospheric subsidence or beginning of downdip motion of slab is accompanied

by penetration of slab-derived fl uids into upwelled mantle, causing melting of depleted harzbur-

gite. (D) Downdip motion of lithosphere signals start of true subduction, which terminates rapid

trench rollback and protoforearc spreading. Forearc mantle cools, and igneous activity retreats

~200 km to what becomes the magmatic arc. Izu-Bonin-Mariana subduction initiation encom-

passed ~7 m.y. for the complete transition from initial seafl oor spreading to normal arc volcanism.

BON—boninite; FAB—forearc basalt; VAB—volcanic arc basalts; LM—lithospheric mantle (gray);

AM—asthenospheric mantle (white). Note tectonic setting of lava sequence shown in Figure 11.

Figure 8. Simplifi ed P-wave veloc-

ity structure beneath Izu forearc

(approximately E-W line at

30°50′E), modifi ed after Kamimura

et al. (2002), with interpreted

lithologies. Note that forearc

crustal thickness decreases from

~11 km near the volcanic front

to 5 km or less near the trench.

Note also signifi cantly lower

P-wave seismic velocity (Vp) in

uppermost mantle beneath the

outer forearc (6.4–6.8 km/s vs.

7.3–8.0 km/s), probably refl ecting

greater serpentinization beneath

the outer forearc.



Mariana forearc, Calvert et al. (2008) concluded that 15-km-thick crust just east of the magmatic front formed during a brief magmatic episode early in the arc’s history. Furthermore, Izu-Bonin-Mariana subduction initiation probably lasted less than 10 m.y.; Ishizuka et al. (2011) estimated that it took ~7 m.y. before a stable magmatic arc was established near the position that it still occupies. Finally, subduction erosion has removed a signifi cant amount of forearc; a width of 1 km removed per million years trans-lates into 240 km3/km of lost crust. All these considerations indicate that Izu-Bonin-Mariana crustal growth related to subduction initiation was signifi cantly greater than 120–180 km3/km.

Such high crustal growth rates and the absence of large, central volcanoes of Eocene age in the Izu-Bonin-Mariana forearc imply that crust was formed by seafl oor spreading, at least during the early, forearc basalts–domi-nated episode. This is consistent with the pres-ence of dense diabase dike swarms at the base of the forearc basalts sequence in the Izu-Bonin-Mariana forearc (Reagan et al., 2010; Ishizuka et al., 2011). Furthermore, the composition of forearc basalts, which is similar to MORB, implies a similar origin by decompression melting beneath a spreading ridge (Plank and Langmuir, 1992). Small central volcanoes may have formed during later, boninitic volcanism, for example, at ODP Site 786B (Lagabrielle et al., 1992), but the thicker, older forearc basalt succession seems to have been emplaced by tectonomagmatic processes akin to seafl oor spreading. There is no clear evidence about the arrangement of spreading ridges that might have existed; Stern (2004) inferred short segments aligned oblique to the evolving convergent plate boundary, although no good magnetic anomaly patterns have yet been identifi ed in any forearc that could be interpreted as spreading fabric.

Given that seafl oor spreading best explains the tectonic environment for Izu-Bonin-Mariana forearc crust formation, the logical conclusion is that a strongly extensional environment existed at that time. There is no evidence that early Izu-Bonin-Mariana subduction was accompa-nied by compression, as would be expected if the new subduction zone was caused by one plate being forced beneath the other (induced subduction initiation of Stern, 2004), although such evidence (uplift-related unconformity, thrust faulting) might have been obliterated by Eocene igneous activity. Such evidence of initial compression without forearc volcanism charac-terizes the Puysegur mini-subduction initiation episode, which serves as an excellent example of induced subduction initiation. The conclusion that a strongly extensional environment accom-panied Izu-Bonin-Mariana subduction initiation

leads logically to the idea that hinged subsid-ence of the older, thicker, and denser Pacifi c plate allowed asthenosphere to well up and fi ll the widening chasm, accompanied by extensive decompression melting (Fig. 8B). The sequence of events summarized in Figure 9 encapsulates the idea of spontaneous subduction initiation (Stern, 2004), including early extension and seafl oor spreading. On the other hand, we can-not rule out the possibility that induced (forced) subduction initiation might also have been asso-ciated with early voluminous igneous activity, due to the likelihood of lithospheric collapse and asthenospheric upwelling, as modeled by Hall et al. (2003).

Does current knowledge about Izu-Bonin-Mariana subduction initiation serve as a useful model for reconstructing how other forearcs form? We cannot be sure because we know so little about their crust, but Izu-Bonin-Mariana serves as a useful analogue for the Tonga-Ker-madec forearc. This crust is exposed only on the island of ‘Eua and was drilled near the trench at ODP Site 841. Arc tholeiite, gabbro, and harz-burgitic peridotite have been dredged from the inner-trench wall, suggesting to Bloomer et al. (1995) that the Tonga forearc is fl oored by crust similar in age and composition to that of Izu-Bonin-Mariana, including Eocene boninite (Crawford et al., 2003). The oldest known rocks in the Tonga-Kermadec forearc are 46–40 Ma arc-type lavas occurring below Upper Middle Eocene limestones on ‘Eua (Ewart and Bryan, 1972; Duncan et al., 1985; Tappin and Balance, 1994). ODP drilling at Site 841 in the Tongan forearc recovered a thick sequence of low-K arc tholeiitic rhyolites (Bloomer et al., 1994), dated by McDougall (1994) at 44 ± 2 Ma. Far-ther north in the Tongan forearc, true low-Ca boninitic rocks and associated backarc basin–type basalts of probable Eocene age have been dredged at depths in excess of 4 km (Falloon et al., 1987) and have yielded ages between 45 and 35 Ma (Bloomer et al., 1998). It is not clear whether or not these boninitic lavas were gen-erated in the same ca. 52 Ma subduction initia-tion event as that responsible for the subduction initiation boninite–refractory-forearc mantle package of the Izu-Bonin-Mariana arc system to the north.

We emphasize that not all subduction zones form by processes outlined here. As outlined by Stern (2004), some subduction zones form as a result of plate-boundary reconfi gurations, for example, as a result of terrane accretion or conti-nental collision. Collision of India with Eurasia is a good example of this, although a new sub-duction zone has not yet formed behind (south of) India, refl ecting the great strength of Indian Ocean lithosphere (Stern, 2004). Collision of

the Ontong-Java Plateau on the north side of the Solomon arc in Miocene time caused a new subduction zone to form to the south (Mann and Taira, 2004), in what may be the best actualistic example of subduction polarity reversal. Stern (2004) characterized this type of subduction initiation, which includes the Puysegur trench example, as refl ecting “induced” nucleation of a subduction zone. Induced nucleation of a sub-duction zone conceptually contrasts with “spon-taneous” nucleation of a subduction zone” such as Izu-Bonin-Mariana may be. Stern (2004) fur-ther suggested that induced nucleation of a sub-duction zone may not be associated with volu-minous forearc igneous activity, such as that forming the Izu-Bonin-Mariana forearc. If so, forearcs formed by induced nucleation of a sub-duction zone may have fundamentally different crustal structures and origins than those formed during spontaneous nucleation of a subduc-tion zone. The generalization that spontaneous nucleation of a subduction zone results in broad forearc magmatism, whereas induced nucleation of a subduction zone does not, may not be true. At least one example of induced nucleation of a subduction zone—Aleutian subduction initia-tion—may be an example of induced nucleation of a subduction zone with attendant forearc igneous activity. Scholl (2007) and Minyuk and Stone (2009) suggested that Aleutian subduction initiation exploited SSW-trending strike-slip faults related to the “tectonic escape” of Alaska, associated with motion along the North Pacifi c Rim Orogenic Stream (Redfi eld et al., 2007). A curved system of long strike-slip faults propa-gated southwestward across the North Pacifi c Rim throughout Cenozoic time, disrupting older subduction zones along the Bering Sea shelf edge. Recently obtained ca. 46 Ma 40Ar/39Ar ages from Aleutian forearc igneous rocks (Jicha et al., 2006; Minyuk and Stone, 2009) provide minimum age constraints for the timing of Aleu-tian subduction initiation, although Aleutian subduction initiation generated middle Eocene magmatic rocks as early as ca. 50 Ma (Scholl, 2007). The effects of tectonic erosion at con-vergent margins must be considered for any thoughtful subduction initiation analysis. Naked forearcs are likely to be trimmed back by sub-duction erosion, at rates that can vary from a few to several kilometers per million years (Clift and Vannucchi, 2004; Scholl and von Huene, 2007). Subduction erosion thus can remove all of the evidence for a magmatic forearc in several tens of millions of years, as may be the case for the Andean forearc.

Keeping such caveats and complications in mind, it seems reasonable to conclude that many forearcs form as a result of voluminous yet ephemeral igneous activity accompanying

STERN ET AL.


subduction initiation. This makes it worthwhile to reconsider the origins of forearc igneous rocks that formed about the same time as sub-duction initiation. For example, the ca. 55 Ma Siletzia terrane of the Oregon and Washington Coast Ranges formed about the same time as Cascadia subduction initiation and is variously interpreted as an accreted oceanic plateau (Dun-can, 1982), produced by the Yellowstone plume head (Pyle et al., 2009), or due to a tear in the subducting slab (Humphreys, 2008). The pos-sibility that Siletzia formed in situ as a forearc magmatic response to Cascadia subduction ini-tiation should also be entertained.

Finally, we should consider the possibil-ity of subduction “re-initiation,” i.e., the case where a subduction zone once existed, then was extinguished, and then resumed at about the same location, because of either induced or spontaneous nucleation of a subduction zone. The Cretaceous and younger evolution of SW Japan serves as an example of this. Subduc-tion of Pacifi c seafl oor beneath NE Japan has been continuous, but subduction of the Philip-pine Sea beneath SW Japan has been episodic. During Cretaceous time, a subduction zone dipped beneath what is now SW Japan, asso-ciated with a robust magmatic arc. Subduction there ceased and was replaced by a transform fault (shear margin) during Paleogene time. This lithospheric weakness was converted into a new subduction zone beginning in latest Oligocene time, with an attendant fl are-up of forearc igne-ous activity (Kimura et al., 2005). Other likely examples of subduction re-initiation include the Paleogene Cascadia system and the Late Juras-sic of California.

To conclude this section, the igneous crust and uppermost mantle of forearcs do not gener-ally represent trapped oceanic lithosphere but in fact typically form during upper-plate spreading associated with subduction initiation (Shervais, 2001; Stern, 2004). There is no doubt that we need to better understand the composition and mode of formation of forearc crust, not only for its own sake but also to better reconstruct events accompanying subduction initiation. Such stud-ies require studying naked forearcs, with all the challenges this entails. In the next section, we propose a complementary strategy that takes advantage of the fact that forearc lithosphere is commonly emplaced (obducted) when buoyant lithosphere—particularly continental crust—on the downgoing plate enters and clogs a sub-duction zone (Wakabayashi and Dilek, 2003). Introduction of buoyant lithosphere disrupts the normal operation of a subduction zone, so that subducted materials are partially regurgitated, and overlying forearc lithosphere is lifted above sea level (Glodny et al., 2005). Consequently,

forearc crust (and accreted sediments) is a key component of orogens. These tracts of obducted forearc lithosphere are known as ophiolites.

OPHIOLITES

Ophiolites are fragments of oceanic litho-sphere that have been tectonically emplaced on land. A complete “Penrose” ophiolite includes tectonized peridotite, gabbro, sheeted dikes, and pillow basalt (Anonymous, 1972), but this idealization is rarely seen because ophiolites are faulted and fragmented during emplacement or because one, or more, unit was never generated. Nevertheless, ophiolites are key petrotectonic indicators, perhaps the single most impor-tant indicator of ancient plate-tectonic activity (Stern, 2005). Ophiolites mark tectonic sutures, indicating both the location of ancient oceans and convergent plate boundaries where buoyant lithosphere was partially subducted, also known as collision zones (Dilek, 2003). As a result, ophiolites are often highly altered and faulted, and much effort and imagination are needed to reconstruct the original crust and uppermost mantle section. Ophiolites are unequivocal evi-dence of seafl oor spreading and have been an

important part of modern geologic thought since the 1960s, but there is a lot of confusion about the tectonic environment in which they formed. Much of this misunderstanding results from a lack of appreciation of the disparate lines of evidence needed to reconstruct ophiolites, espe-cially structural geology, igneous geochemistry, and marine geology. Ophiolites were originally thought to form at mid-ocean ridges, but we now understand that most sediments and all crust on the downgoing plate are subducted. If the down-going plate includes very thick (>1 km; Clift and Vannucchi, 2004) sediments, some of this may be scraped off, and sometimes seamounts may be transferred from the subducting to the over-riding plate, but normal oceanic crust itself has never been demonstrated to be transferred from downgoing to overriding plate (i.e., by seismic-refl ection profi ling) at any modern convergent margin (Fig. 10C). Changes in plate motion might trap mid-ocean-ridge crust in what may ultimately become a forearc (e.g., Macquarie Island; Varne et al., 2000), but this is likely a very unusual tectonic scenario.

In the 1970s, geoscientists (beginning with Miyashiro, 1973) began to recognize the simi-larity of some ophiolite lava compositions to

Figure 10. Tectonic cartoon illustrating the relative feasibility of emplacing oceanic lithospheres

created in forearc, backarc basin, and mid-ocean-ridge settings. (A) It is relatively easy to emplace

forearc lithosphere. Subduction of buoyant material commonly leads to failure of subduction zone,

and isostatic rebound of buoyant material emplaces ophiolite. (B) It is diffi cult to emplace backarc

basin oceanic lithosphere. Compression and shortening across the arc will lead to uplift of the arc.

(C) It is almost impossible to emplace true mid-ocean-ridge basalt (MORB) crust at a convergent

plate boundary. Sediments and fragments of seamounts may be scraped off the downgoing plate,

but subducting MORB-type lithosphere is nowhere known to be transferred from downgoing to

overriding plate (modifi ed after Stern, 2004).

Backarc basinlithosphere

Ophiolite

Forearc lithosphere

MORB lithosphere

Forearc ophiolite: Easy to emplace Backarc ophiolite: Difficult to emplace

MORB ophiolite: Almost impossible to emplace

Ophiolite

Buoyant crust

Subducting plate

Suture

A B

C



those of convergent margins, leading to the concept of the “suprasubduction zone” ophio-lite (Pearce et al., 1984). This created tension between conclusions based on structural stud-ies, which indicated ophiolites form by seafl oor spreading, and those based on geochemistry, which indicated ophiolites form at convergent margins. This tension was temporarily rec-onciled by the idea that suprasubduction zone ophiolites formed in backarc basins (Dewey, 1976; Pearce, 2003). This reconciled structural evidence of seafl oor spreading with geochemi-cal evidence for convergent margin setting, but it is diffi cult to see how a backarc basin, ~200 km from the convergent plate boundary, would be emplaced (Fig. 10B). This logic train is further derailed for some ophiolites interpreted as fossil backarc basins because the associated arc and forearc, which are typically ~200 km wide and ~30 km thick, are usually not identifi ed (e.g., Semail ophiolite interpretation of Godard et al., 2003). Furthermore, there are no known Ceno-zoic examples of backarc basin closure (Stern, 2004), which is required to emplace a backarc basin ophiolite.

A simple and elegant solution to the ophio-lite emplacement problem is that whatever crust comprises the forearc is most likely to be emplaced during plate collision. This process is often referred to as “obduction,” although this term originally described how oceanic crust on the downgoing plate was thrust over the con-vergent margin (Coleman, 1971). Forearc litho-sphere is optimally situated for obduction (Wak-abayashi and Dilek, 2003). It is straightforward to envisage emplacement of forearc lithosphere above the same subduction system in which it was generated due to isostatic uplift following partial subduction of buoyant crust that jams the subduction zone (Fig. 10A). This process—akin to sliding a spatula under a pancake or fried egg to lift it—is also most likely to yield the least-disrupted ophiolites, and those most likely to approximate the Penrose ophiolite ideal, such as Troodos (Cyprus) and Semail (Oman). Dur-ing the early days following the plate-tectonic revolution, when it was thought that forearc crust was relict, trapped oceanic crust, this provided an attractive way to emplace MORB-type ophio-lites. As discussed previously herein, there is lit-tle evidence to support the idea that forearcs are composed of trapped oceanic crust that existed in the region prior to the formation of a new sub-duction zone. It is increasingly clear from the simple perspective of plausible emplacement mechanisms that forearcs are the most likely source of ophiolites (Casey and Dewey, 1984; Milson, 2003; Metcalf and Shervais, 2008).

Because most ophiolites are very incom-plete and rarely satisfy the Penrose defi nition,

we propose an alternative defi nition that bet-ter captures the important elements of what are called ophiolites in the literature. An ophiolite must consist of a signifi cant proportion of harz-burgitic peridotite (depleted mantle) and pil-lowed basalt. Gabbro and sheeted dikes may be missing, but there should be associated deep-sea sediments. These units should be exposed above sea level. This defi nition is more consis-tent with common usage of the term ophiolite than is the Penrose defi nition.

Based on emplacement mechanisms, Waka-bayashi and Dilek (2003) distinguished four types of ophiolites: (1) “Tethyan” ophiolites, emplaced over passive continental margins; (2) “Cordille-ran” ophiolites, emplaced over subduction com-plexes; (3) “ridge-trench intersection” ophiolites emplaced through complex processes resulting from the interaction between a spreading ridge and a subduction zone; and (4) the unique Mac-quarie Island ophiolite, which was subaerially exposed as a result of a change in plate-boundary confi guration along a mid-ocean-ridge system. The fi rst two types represent the overwhelming majority of ophiolites, with the second two types comprising miniscule proportions. Tethyan- and Cordilleran-type ophiolites refl ect the fundamen-tally different nature of the subducted oceanic realms, with Tethyan ophiolites subducting rela-tively narrow oceans before colliding with conti-nental fragments and Cordilleran ophiolites gen-erally subducting Pacifi c seafl oor, which contains no continental fragments.

Given that most ophiolites originate in forearcs, is it possible to independently deter-mine whether or not a given ophiolite formed during subduction initiation or as a result of some other process, as outlined herein? Such an assessment would be very useful because of the diffi culties involved in studying submerged forearc igneous rocks and reconstructing sub-duction initiation processes. If ophiolites could be so linked, it could pay huge dividends in advancing our understanding of the fundamen-tal tectonic province of forearc and processes of subduction initiation. The next section summa-rizes a chemostratigraphic approach for evalu-ating whether or not a given ophiolite formed during subduction initiation.

THE SUBDUCTION INITIATION RULE

The subduction initiation rule (Whattam and Stern, 2011) predicts that ophiolites that form as a result of subduction initiation processes consist of a sequence of igneous rocks formed by a magma source that changed progressively in composition by the combined effects of melt depletion and subduction-related metasoma-tism. Magmas erupted during subduction initia-

tion progress from early decompression melts of unmodifi ed fertile (lherzolitic) mantle to yield forearc basalts to younger hydrous fl ux melts of depleted (harzburgitic) mantle that has been strongly modifi ed by subduction-related fl uids to yield late high-Mg andesites and boninitic lavas. If the subduction initiation process con-tinues long enough to generate steady-state sub-duction with downwelling mantle overlying the sinking plate, then normal arc volcanics will cap the subduction initiation sequence as the locus of magmatism retreats from the trench. This magmagenetic evolution is portrayed in Fig-ures 7 and 11, which encapsulate the subduction initiation tectonic evolution shown in Figure 9. Whattam and Stern (2011) outlined the argu-ments and implications of the subduction initia-tion rule and proposed that ophiolites showing this progression be termed “subduction initia-tion rule ophiolites.”

It has long been recognized that some ophiolites contain igneous rocks with strong chemical affi nities to both mid-ocean-ridge and arc basalts. Distinguishing compositions of MORB-like lavas include >1% TiO

2, vari-

able depletion in LREEs, absence of HFSEs (e.g., Nb, Ta), depletions on spider (primitive mantle– or N-MORB–normalized) diagrams, La/Nb <1, etc.; we regard these early MORB-like sequences as forearc basalts. In contrast, volcanic arc–like basalts have lower TiO

2, are

enriched in fl uid-mobile elements (e.g., LILEs and LREEs), and have strong HFSE depletions relative to LREE (e.g., La/Nb >1, etc.; Pearce, 2003). Recognition of both arc-like and MORB-like compositions in some ophiolites has encouraged some workers to infer formation in a backarc basin (e.g., Beccaluva et al., 2004) or a complex, multistage tectonic history, for exam-ple, eruption of the two suites in discrete tec-tonic environments (e.g., Saccani and Photiades, 2004; Godard et al., 2003). The reasons against formation as a backarc basin are outlined in the previous section. The conclusion here that the basaltic sections in ophiolites are forearc basalts (Reagan et al., 2010) is supported by the obser-vation that most Izu-Bonin-Mariana forearc basalts have lower Ti/V ratios than MORB, which probably results from the enhanced melt-ing that commonly occurs in nascent subduction settings. The presence of transitional lavas with compositions that progress from forearc basalts to boninite with time at DSDP Sites 458 and 459 also supports this contention. Some ophiolites may refl ect complex tectonic histories, but this cannot serve as the general explanation for sub-duction initiation rule ophiolites. Furthermore, any inference of complex tectonic histories is inconsistent with the absence of unconformi-ties or signifi cant sedimentary horizons between

STERN ET AL.


ophiolite lavas with MORB-like versus volca-nic arc–like basalt compositions. Such inter-ruptions might be diffi cult to identify because of ophiolite deformation, but such an important and distinctive feature should sometimes be recognized if this interpretation is generally cor-rect. Conversely, the interpretation of a single, rapidly evolving magmagenetic environment is favored because such a hiatus is generally not observed. In addition, lavas with compositions that are transitional between forearc basalts and boninites have been observed in some ophiolites (e.g., Dilek and Furnes, 2009). As a result, ophi-olite compositional variability is increasingly ascribed to progressive depletion and metaso-matism of the mantle source as the ophiolite magmatic crust forms (e.g., Shervais, 2001; Beccaluva at al., 2005; Dilek and Furnes, 2009). Ophiolite lava compositions thus are increas-

ingly interpreted to be products of magmatism in a single (suprasubduction zone) tectonic set-ting, albeit one that changed rapidly.

It is worth emphasizing that it is generally diffi cult to recognize a magmatic stratigraphy in ophiolites, because they are so often jumbled by faulting, and because lavas and associated intru-sions with different chemistries appear similar in the fi eld. We recognize four ophiolites that have been studied in suffi cient detail to recon-struct their magmatic stratigraphies: Mirdita, Pindos, Troodos, and Semail (as summarized by Whattam and Stern, 2011). There are other examples of subduction initiation rule ophiolites that could also be considered, for example, the 485–489 Ma ophiolite belt that can be traced >1000 km from Newfoundland down into the Taconic suture of New York (Bédard et al., 1998; Schroetter et al., 2003; Pagé et al., 2009).

The four Tethyan ophiolites mentioned here show subduction initiation rule relationships, with lower lavas being more MORB-like and upper lavas being more arc-like, as Whattam and Stern (2011) showed. Troodos and Semail are at either end of the 3000-km-long ophiolite belt that marks the Late Cretaceous forearc of SW Asia, including the inner and outer ophio-lite belts of Zagros, Iran (Moghadam and Stern, 2011). The magmatic chemostratigraphies of the many Zagros ophiolites have not been worked out, but these ophiolites in many cases have strong compositional affi nities in some respect to arc magmas, for example, with respect to La/Nb (arc lavas usually have La/Nb >1.4 accord-ing to Condie, 1999). It is very likely that we will be able to work out magmatic stratigraphies in more ophiolites around the world, but this will require thoughtful integrated structural and petrologic studies.

DISCUSSION AND CONCLUSIONS

Subduction initiation rule ophiolites provide wonderful opportunities to better understand the crust and upper mantle of magmatic forearcs and the ways in which new subduction zones form. Identifying and studying subduction ini-tiation rule ophiolites will provide easy access to forearc lithosphere samples and structures, allowing many scientifi c perspectives to be engaged cheaply and easily. Where ophiolites can be fi rmly linked to forearcs through applica-tion of the subduction initiation rule and other approaches, geodynamic models for subduc-tion initiation can be more easily developed and tested. Because subduction initiation rule ophiolites are fossil forearcs, they often can be traced across strike for tens of kilometers and along strike for hundreds of kilometers. Ero-sion of deformed subduction initiation rule ophiolites exposes various levels, allowing four-dimensional reconstructions of timing as well as vertical, across-strike, and along-strike mag-matic variations. Sampling for geochronology is easier than for in situ forearcs because of this exposure, and relations of such dated samples to the surrounding rocks and fabrics provide con-text for interpreting the ages. The multiple per-spectives and levels of detail allowed by these approaches mean that many aspects of forearc crust structure and subduction initiation can be understood much better by indirect study of sub-duction initiation rule ophiolites than by study-ing forearcs directly, as can be seen by mentally comparing the experiences captured in Figure 1.

Even as we recognize that studies of sub-duction initiation rule ophiolites are essential for understanding forearcs, studies of in situ forearc crust need to move forward. We are only

Oldest

Youngest

Tholeiitic

LREE-depleted

Higher ΣREE

Higher TiO , Y, Zr

Higher Ti/V

Lower Cr/Y

Calc-alkaline

Sometimes boninite

Mostly LREE-depleted

Lower ΣREE

Higher TiO , Y, Zr

Lower Ti/V

Higher Cr/Y

Lherzolitic mantle sourceUnaffected by fluids from subducted crust /sediments

Harzburgitic mantle sourceModified by fluids from subducted crust /sediments

Lava Mantle Age composition composition

The Subduction Initiation Rule

1–

3 k

m

2

2

Figure 11. The subduction initiation rule, simplifi ed after Whattam and Stern (2011). Ophiolites

(and forearc crust) that form as a result of subduction initiation preserve systematic variations in

basalt compositions (left), from mid-ocean-ridge basalt (MORB)–like at the base to volcanic arc)–

like basalt (VAB), even boninitic, at the top. This refl ects changes in the source mantle as a new

subduction zone starts, beginning with adiabatic upwelling of asthenosphere that is not infl u-

enced by subduction-related fl uids (Fig. 9B) to form MORB-like basalts (forearc basalts) by seafl oor

spreading. Fluids from the sinking lithosphere become increasingly important with time, eventu-

ally reaching and metasomatically re-enriching the increasingly depleted mantle source of melts

(Fig. 9C). Asthenospheric upwelling diminishes in importance with time and is ultimately replaced

by induced convection as sinking lithosphere begins downdip motion, and true subduction begins

(Fig. 9D). This isolates the forearc mantle wedge, leading to establishment of localized magmatic

arc behind a cold, dead forearc. LREE—light rare earth element; ∑REE— total rare earth element

concentrations.



beginning to understand the igneous crust of a single forearc in any detail—that of Izu-Bonin-Mariana—and this example seems to be unusual in terms of the abundance of early-arc boninites. We need to continue to test and refi ne what we know about forearc magmatic evolution. It is especially important that we drill and sample through the magmatic stratigraphy of an igneous forearc in order to better understand the sequence and timing of magmas at one location, allowing us to answer questions such as: What are the relative proportions of forearc basalts versus vol-canic arc–like basalt/boninitic lavas? How does forearc basalt magmatism transition to volcanic arc–like basalt magmatism? Is it gradational or abrupt? How long does forearc volcanism last? Answering such questions for an in situ forearc and comparing these answers with those for sub-duction initiation rule ophiolites will be key tests of the ideas presented in this paper. Furthermore, there are aspects of forearc structure and evolu-tion that cannot be understood without studying extant forearcs. For example, active serpentine mud volcanoes in the Mariana forearc serve as actualistic models for sedimentary serpentinite deposits associated with some forearcs (Fryer et al., 1995; Fryer, 2002). Knowing that these mud volcanoes exist has stimulated the search for ancient serpentine mudfl ows (e.g., Teklay, 2006). Another example is the problem of tec-tonic erosion, which can remove a few kilome-ters of forearc crust in a few million years. Stud-ies of active systems are generally aware that some crust is missing, whereas studies of ophiol-ites rarely consider this complication.

Another point worth emphasizing is that there may be a wide range of subduction ini-tiation magmatic products. Our models now are heavily biased toward the Izu-Bonin-Mar-iana convergent margin, which has abundant boninitic lavas and serpentinite mud volcanoes. The Izu-Bonin-Mariana forearc may serve as a useful example of intra-oceanic subduction initiation, but it could be less useful as a model for continental margin subduction initiation. Subcontinental mantle lithosphere is compo-sitionally distinct from that beneath the ocean basins (Bonatti and Michael, 1989), and conti-nental and oceanic mantle may melt to different extents during subduction initiation. Because there are likely to be signifi cant differences in magmatic products from that shown by Izu-Bonin-Mariana and ophiolites, we should keep an open mind when considering the origin of any magmatic forearc built about the time that subduction began at a given convergent mar-gin. The Siletzia terrane of coastal Oregon and Washington (Duncan, 1982) could be an exam-ple of continental margin subduction initiation where the magmatic sequence is somewhat dif-

ferent than that seen for Izu-Bonin-Mariana, for example, by having upper alkalic lavas instead of volcanic arc–like basalt overlying tholeiites.

Above all, combined studies of subduction initiation rule ophiolites and forearc crust need to be integrated with geodynamic modeling to learn about the ways in which new subduction zones form. Geodynamic modeling of subduc-tion initiation needs to be fi rmly tethered in real-ity, and the more that such models attempt to explain the rocks making up a real forearc and ophiolite, the more rapidly our understanding of this fundamental Earth processes will advance.

ACKNOWLEDGMENTS

We thank Dave Scholl for edifying comments on accretionary prisms and Steve Graham for the photo of a Franciscan chert exposure. Thought-ful reviews by Rod Metcalf, Jean Bédard, and editors John Shervais and John Goodge are much appreciated. This manuscript was written while Stern enjoyed a Blaustein Fellowship at Stanford University. Stern’s research on intra-oceanic arcs has been supported by the National Science Foundation, most recently by grant OCE-0961352. Reagan acknowledges research support from National Science Foundation grant EAR-0840862. This is University of Texas at Dallas Geosciences contribution 1228.

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