The forearc lies between the trench and the magmatic front and is 166 ± 60 km wide [Gill, 1981]. A first-order distinction can be made between forearcs that broaden by addition and imbrication of material along the trench (accretionary margins) and those that do not (Figure 16). This depends on the thickness of sediment being subducted. If greater than 400–1000 m, sediments will be scraped off the downgoing plate and transferred to the overriding plate to form an accretionary prism [Dahlen, 1990; von Huene and Scholl, 1991; Le Pichon et al., 1993]. The total length of convergent plate margins is about equally divided between accretionary and nonaccretionary types [von Huene and Scholl, 1993]. About 80% of sediments that arrive at trenches are not accreted [von Huene and Scholl, 1991].
Figure 16. End-member forearc types: (a) accretionary forearc (modified after Dickinson ) and (b) nonaccretionary forearc (P. Fryer, personal communication, 2001). Note that the abundance of sediments associated with accretionary forearcs is manifested as an accretionary prism and as a thick forearc basin and that the relative lack of sediments leaves the nonaccretionary forearc exposed.
Download figure to PowerPoint
 Accretionary forearcs form where sediment supply is high, typically adjacent to or near continents. Accretionary forearcs grow by frontal accretion, as sediment is “bulldozed” off the subducting plate to form an accretionary wedge or prism. These are highly deformed and composed of arcward dipping, thrust-bound packets with emplacement ages that become younger toward the trench [Underwood and Moore, 1995]. The forearc can also be thickened by underplating of subducted sediments in the “subduction channel” (Figure 16a) [Cloos and Shreve, 1988a]. The high sedimentation rates associated with accretionary forearcs also make these important depocenters, with robust forearc basins forming between the accretionary prism and the magmatic arc (Figure 16a) [Dickinson, 1995].
 Nonaccretionary forearcs form where sediment supply is low, typically distant from continents, and thus they also have poorly developed forearc basins. Because the accretionary prism and forearc basin are missing, the igneous infrastructure of the forearc is exposed. Nonaccretionary forearcs provide unique insights into how subduction zones begin and their early history, the significance of ophiolites, and the nature of fluids released from relatively shallow parts of subduction zones.
 One of the major uncertainties in nonaccretionary forearcs concerns how much of the forearc is removed by “subduction erosion” or “tectonic erosion,” whereby crust and mantle of the upper plate is carried down with the subducting plate. Removal of significant volumes of the forearcs of NE Japan, Central America, South America, the Marianas, and Tonga has been inferred. This is thought to have occurred as material was removed from the base of the upper plate (basal erosion) and by slumping of material into the trench, which then was carried down (frontal erosion). Mechanisms for basal tectonic erosion include “rasping” at the base of the forearc by seafloor which has been roughened by horst-and-graben faulting developed between the forearc bulge and the trench [Hilde, 1983]. Frontal erosion is inferred from disrupted topography at the base of the inner trench wall and is most evident in the wake of subducting seamounts [von Huene and Scholl, 1991].
 Forearc basement is relatively cool and stable and so preserves a record of how and when the associated subduction zone began; this record is accessible in nonaccretionary forearcs. Forearcs of the Izu-Bonin-Mariana and Tonga-Kermadec arc systems contain an especially clear record of the early history of these subduction zones. Earlier ideas that forearcs are largely composed of “trapped” oceanic crust that formed before the pertinent subduction zone [Dickinson and Seely, 1979] are not generally supported by recent studies. Forearc crust generally seems to have been formed by seafloor spreading about the same time as its subduction zone was initiated [Bloomer et al., 1995], although a minor amount of trapped oceanic crust has been identified [DeBari et al., 1999]. This indicates that subduction initiation is associated with early extension, not shortening [Stern and Bloomer, 1992]. Most ophiolites appear to be samples of forearc basement, formed at the time of subduction initiation and emplaced when subduction zones are terminated by collision [Bloomer et al., 1995; Shervais, 2001].
 The Mariana forearc has a number of serpentine mud volcanoes, up to 2 km tall and 30 km in diameter, which “erupt” serpentine mud [Stern and Smoot, 1998; Fryer et al., 2000]; this is the only convergent margin where active serpentine mud volcanism is known. Flows from these mud volcanoes carry abundant cobbles and boulders of serpentinized peridotite and minor metabasalt and metagabbro. Refractory peridotites (mostly harzburgite and minor dunite) are most common [Ishii et al., 1992]. Harzburgite spinels are generally more Cr-rich than those found in MORB-type abyssal peridotites, indicating that these are residues left after extensive melting. Parkinson and Pearce  interpret the trace element signature of some of these harzburgites as residual MORB mantle (15–20% fractional melting) subsequently modified by interaction with boninitic melt. These peridotites are similar to those recovered from the Mariana inner trench wall.
 Mafic crustal blocks (both arc and MORB affinities) are ∼10% as common as peridotite blocks in the serpentine mud. Low-temperature metamorphism affects most samples, ranging from zeolite to lower greenschist facies. Maekawa et al.  documented clasts containing blueschist-facies mineralogy (lawsonite, aragonite, sodic pyroxene, and blue amphibole) and used these to estimate metamorphic conditions of 150°–250°C and 5–6 kbar in the subducted plate. This is the first documented link between an active subduction zone and blueschist-facies metamorphism.
 Vent chimneys composed of aragonite, calcite, and hydrated Mg-silicate are found near summits of active serpentine mud volcanoes. Compared to ambient seawater, fluids emanating from vents in the summit area are slightly cooler and have higher pH and alkalinity. The vent waters are enriched in methane, silica, and H2S [Fryer et al., 1990]. Pore waters from the cores drilled near the summit of Conical Seamount are some of the most unusual ever sampled in oceanic sediments, containing less than half of the chloride and bromide concentrations of seawater, pH up to 12.6, and methane along with ethane and propane [Mottl, 1992]. Relative to seawater, these fluids are enriched in alkalinity, sulfate, K, Rb, and B and are depleted in Li, Mg, Ca, and Sr. Mottl  concluded that these fluids originated from the downgoing slab, 30 km beneath the seafloor. Sampling fluids from serpentine mud volcanoes built at different heights above the Mariana subduction zone suggest that systematic variations in fluid chemistry manifest progressive decarbonation reactions in the downgoing slab [Fryer et al., 1999].
 Two issues relating to forearc lithosphere await resolution. The first concerns the definition of lithospheric mantle, which, on the basis of observed low heat flow, clearly underlies normal forearcs. Lithospheric mantle is normally thought to be colder, denser, and stronger than asthenosphere [Anderson, 1995], but beneath forearcs, extensive hydration and serpentinization may make the “lithosphere” colder but weaker and less dense than asthenosphere. In this case, how is forearc lithosphere to be defined, and how can the lithosphere-asthenosphere boundary beneath forearcs be geophysically identified? The second issue concerns the geometry of the boundary between convecting asthenosphere and forearc “lithosphere.” All models for convective flow in the mantle above the subduction zone require corner flow, where asthenosphere descends. Isotherms in convecting asthenosphere parallel the mantle flow, to a first approximation, so the lithosphere-asthenosphere boundary may be more vertical than horizontal, as is normally the case [e.g., see Takahashi et al., 1998, Figures 1b and 9].
6.2. Magmatic Arc
 There are three principal tectonic settings where melted mantle moves to Earth's surface: island arcs, mid-ocean ridges, and hot spots. Each is caused and manifested differently. Mid-ocean ridges are line source loci of volcanism reflecting two-dimensional (2-D) upwelling and decompression melting of shallow mantle. Hot spots are point source volcanic loci manifesting isolated regions in the deep mantle, which melt because of the intrinsic properties of the region itself. To a first approximation, magmatic arcs are linear arrays of point sources, caused by Rayleigh-Taylor instabilities in the mantle wedge above subduction zones [Marsh, 1979]. As discussed in section 5.3, this melt zone results when hydrous fluids released from subducted materials drastically lower the melting temperature of the overlying asthenospheric mantle.
 Magmatic arcs are actually more like ribbons than lines of volcanoes, averaging 97 km wide [d'Ars et al., 1995]. The magmatic front marks the boundary between low heat flow in the forearc and high heat flow beneath the magmatic arc and the back arc region (Figure 1b). Igneous activity is concentrated nearest the magmatic front and diminishes with greater distance from the trench. The magmatic front lies 124 ± 38 km above the inclined seismic zone [Gill, 1981], and this relationship does not vary systematically with any subduction variable, such as convergence rate or age of lithosphere being subducted. This relationship is probably controlled by the depth at which the slab must lie in order to allow sufficiently hot asthenosphere to be present in the hot corner such that addition of water from the subducted slab leads to melting (Figure 11). There does not seem to be a characteristic spacing between arc volcanoes [d'Ars et al., 1995]. Volcano spacing may be controlled by the thickness of the upper plate lithosphere [Vogt, 1974], gravitational instability in the melt source [Marsh, 1979], or distance between asthenospheric “hot fingers” advected into the mantle wedge [Tamura et al., 2002].
 Arc magmas are generally fractionated, porphyritic, and wet [Perfit et al., 1980; Ewart, 1982; Tatsumi and Eggins, 1995], especially when compared to mid-ocean ridge or hot spot magmas. These three observations are related. Arcs are the only one of the three great magmatic regimes where the region in the mantle that is melting and the overlying crust and lithosphere stay relatively fixed. Consequently, crustal thickening is characteristic of arcs. The thickening rate of intraoceanic arc crust is ∼350 m Myr−1 for the Izu arc. Thicker crust is more difficult for mafic magmas to rise through, especially if it is low-density continental crust; consequently, arc magmas tend to stagnate in the crust, where they fractionate and assimilate. The tendency of arc magmas to fractionate is reinforced by loss of magmatic water at crystal pressures, which results in crystallization even without cooling.
 Olivine, pyroxene, hornblende, and especially plagioclase are the typical phenocrysts of arc lavas. Arc lavas are generally not in equilibrium with mantle peridotite but have experienced significant low-pressure fractionation, usually within the crust. The porphyritic nature of arc lavas indicates that these are mostly mixtures of melt and crystals, so that analyzing bulk lava samples may give misleading information about how these melts evolve. Fortunately, trapped melts in the form of glass inclusions are common in arc phenocrysts, and the composition of these glasses gives more reliable information about melt evolution [Lee and Stern, 1998; Kent and Elliott, 2002].
 Arc lavas are dominantly silica-oversaturated and subalkaline and are further subdivided into volumetrically predominant calc-alkaline and tholeiitic suites and less common boninitic and shoshonitic suites. A useful distinction is that arc tholeiites plot in the low-K portion of Figure 17, calc-alkaline lavas lie within the medium and high-K fields, and shoshonitic lavas are yet more enriched in potassium and other incompatible trace elements. Boninites are a type of high-magnesium andesite requiring unusual conditions of hot mantle and abundant water. These conditions are likely to occur only during the early stages in the evolution of an arc, as large amounts of water are injected into normal asthenosphere. Arc magmas may evolve from boninitic to tholeiitic to calc-alkaline to shoshonitic over the 40–100 million year lifetime of a typical arc [Hawkins et al., 1984; Jolly et al., 2001].
Figure 17. Potassium-silica diagram for representative arcs. Dashed line defines boundary between shoshonitic and calc-alkaline and tholeiitic suites (CATS). Izu-Bonin and Mariana arcs are exemplary of intraoceanic arcs (fields from Stern et al. , note that volumetrically subordinate Mariana shoshonites [Sun and Stern, 2001] are omitted). Field for Andes, 16°S–26°S encompasses most of 606 Plio-Pleistocene and younger samples from the Central Volcanic Zone (CVZ) (G. Wörner, personal communication, 2002). Abbreviations are as follows: M, typical MORB from Table 2; B, back arc basin basalt from Table 2; I, mean composition of Izu-Bonin arc samples; MA, mean composition of Mariana arc samples; CC, bulk continental crust from Table 2; G, GLOSS from Table 2; A, mean composition of Andes CVZ lavas; and UC, composition of upper continental crust [from McLennan, 2001]. Dark shading encompasses mean and typical compositions of the “oceanic suite” (MORB, back arc basin basalt, and mean Mariana and Izu-Bonin lavas), and light shading encompasses mean and typical compositions of the “continental suite” (bulk continental crust, GLOSS, upper continental crust, and mean Andean dacite).
Download figure to PowerPoint
 Arc magmas contain up to 6 wt % H2O, compared to generally <0.4% H2O for MORB and <1.0% H2O for hot spot tholeiites such as Kilauea [Johnson et al., 1994]. Back arc basin basalts (discussed in section 7) are also wetter than MORB and oceanic island basalts but generally contain less water than arc lavas (Figure 18). The main effects of H2O on crystallization of mafic melts are to decrease melt liquidus temperature and to suppress plagioclase crystallization relative to olivine and clinopyroxene [Danyushevsky, 2001]. The water- and silica-rich nature of arc magmas results in violent eruptions and debris flows. This eruption style, coupled with the location of many arc volcanoes near population centers and air traffic routes, leads to these eruptions being by far the most dangerous on the planet [Tilling, 1996]. Additional water is lost by slow degassing of magma at moderate pressures in the crust, so that arc lavas are quite dry. Fortunately, microanalytical techniques permit analysis of glass inclusions trapped in arc lava phenocrysts, yielding better estimates of how much water was contained in the magma prior to degassing [Sisson and Layne, 1993; Newman et al., 2000].
 The CO2 budget for magmatic arcs is problematic. Sano and Williams  and Marty and Tolstikhim  conclude that arc volcanoes release an order of magnitude more CO2 than MORB or hot spot lavas, but glass inclusions in phenocrysts from arc lavas contain vanishingly little CO2. Part of this discrepancy may be due to the very low solubility of CO2 in CO2-H2O mixtures, even at pressures corresponding to the middle to lower crust [Newman et al., 2000]. Lower crustal degassing of CO2 could be proved by documenting exposed sections of arc lower crust that have been extensively modified by pervasive CO2 streaming, but such examples are unknown to the author. Extensive degassing of CO2 from arc magmatic systems is also difficult to reconcile with the results of thermodynamic modeling, which predicts little carbonate breakdown in subduction zones (section 4.2). The issue of how much CO2 is released from magmatic arcs, and at what depth, is worthy of further investigation.
 Arc lavas are characteristically more fractionated than those erupted from hot spots or mid-ocean ridges, but those from Andean-type arcs are more fractionated than those from intraoceanic arcs (Figure 17). Lavas from the Andes are rich in silica and potassium, mostly plotting in the field of high-K andesites and dacites, whereas lavas from the intraoceanic Mariana and Izu-Bonin arcs are mostly medium-K or low-K basaltic andesites. Mean compositions for the intraoceanic arc lavas are more similar to that of typical oceanic crust (MORB), whereas the Andean lava is very close to the composition of upper continental crust. This allows for arc lavas to be subdivided into “continental” and “oceanic” suites (Figure 17). These two great suites of arc lavas can also be distinguished on spidergrams (Figure 14) because the continental suite is more enriched in incompatible elements. These enrichments reflect the thicker, more felsic nature of continental crust underlying Andean-type arcs. The relatively low density of continental crust causes mantle-derived basaltic melts to stagnate within it or at its base [Herzberg et al., 1983], encouraging fractionation and assimilation. Continental crust melts at a relatively low temperature, encouraging melting of crust by mantle-derived melt [Hildreth and Moorbath, 1988]. Felsic melts, including the great calc-alkaline batholiths such as those of the Mesozoic circum-Pacific, may result from melting the crust or from combined assimilation and fractional crystallization in crustal magma reservoirs [DePaolo, 1981]. In contrast, the thinner, mafic crust of juvenile arcs is denser and melts at higher temperatures, discouraging stagnation, fractionation, and crustal melting [Pearcy et al., 1990; Miller and Christensen, 1994], but even juvenile arcs like Izu-Bonin may have a midcrustal tonalite layer which formed as a result of anatexis of mafic arc crust (Figure 19) [Kawate and Arima, 1998].
Figure 19. Magmatic arc complexities. Only an idealized section through an intraoceanic arc is shown; similar processes are expected beneath Andean-type arcs. Note that the asthenosphere is shown extending up to the base of the crust; delamination or negative diapirism is shown, with blocks of the lower crust sinking into and being abraded by convecting mantle. Regions where degassing of CO2 and H2O is expected are also shown.
Download figure to PowerPoint
 Thick, compositionally zoned and layered batholiths and eruptions of zoned magma bodies provide further evidence that fractionation in the crust is an important part of arc melt evolution. The common association of craters and calderas at the summits of arc volcanoes suggests that the substrate is weak, as would be expected for shallow magma chambers. In spite of these indications, there is no seismic tomographic evidence for an extensive magma chamber beneath any active arc volcano. It is instructive to recall that mid-ocean ridges were first thought, on the basis of petrologic arguments and the ophiolite model, to be associated with cavernous magma chambers. Magma reservoirs beneath spreading ridges are now known to be more like thin ribbons of magma, even for the most magmatically robust ridges [Solomon and Toomey, 1992]. On the basis of this experience and the lack of geophysical evidence to date, models calling for the common association of large magma chambers beneath arc volcanoes should be regarded with suspicion.
 Although there is general agreement that juvenile continental crust is mostly formed at magmatic arcs [Reymer and Schubert, 1984], we have much to learn about how this happens. Crustal thickening beneath the magmatic arc can readily be accomplished by overplating of lavas and underplating and “interplating” of plutons and cumulates. Underplating is largely accomplished by formation of thick sequences of mafic and ultramafic cumulates, which must exist to complement abundant fractionated arc lavas [DeBari, 1997]. For the Aleutian arc, Kay and Kay  estimated that 2/3–3/4 as much upper crust is added as lower crust, mostly as cumulates. Intrusions in the lower crust or near the Moho can also thicken the crust. At the same time that arcs thicken by piling lavas on top and plastering cumulates and sill complexes below, existing crust is continuously being reprocessed. A high-quality profile through Izu arc crust reveals a thick layer characterized by Vp ∼ 6.0 km s−1, which has been interpreted as tonalitic middle crust (Figure 15a) [Suyehiro et al., 1996]. Similar tonalites are exposed in the Tanzawa Mountains of Japan [Kawate and Arima, 1998]. These are interpreted as melts of amphibolites in the middle crust [Nakajima and Arima, 1998] that have been exposed by uplift due to collision between the Izu arc and Honshu. Intriguingly, a similar “tonalitic” layer was not found for Aleutian arc crust (Figure 15b) [Holbrook et al., 1999], although tonalitic and granodiorite plutonic rocks outcrop along the Aleutians [Kay and Kay, 1994]. Midcrustal low-velocity zones are known from beneath the Andean arc [Yuan et al., 2000; Haberland and Rietbock, 2001]; these are likely zones where mantle-derived mafic melts fractionate and assimilate older continental crust [Hildreth and Moorbath, 1988]. Midcrustal anatexis should also provide dense, mafic lower crustal material as residues.
 The discrepancy between the basaltic composition of arc melts derived from the mantle compared to the andesitic bulk composition of the continental crust remains as an outstanding problem [Kelemen, 1995; Rudnick and Fountain, 1995]. This can be seen by comparing the mean compositions of intraoceanic arc lavas with the bulk composition of continental crust (Table 2 and Figure 17). An attractive possibility for thick arcs is that mafic cumulates, which belong to the crust in a petrologic sense, contain a significant proportion of dense dunites, pyroxenites, and garnet-bearing eclogitic rocks, which have seismic velocities similar to mantle peridotite. Similar cumulate ultramafics and garnet-bearing “eclogites” at the base of the Mesozoic magmatic arc are documented for the Sierra Nevada batholith of California [Ducea and Saleeby, 1998]. If these dense cumulates are considered to be part of the mantle, as is likely if crustal thickness is inferred from seismic velocities, then the composition of bulk arc crust would be calculated to be more felsic than the bulk composition of the melts rising from the mantle to the crust.
 Delamination or “drip” of dense mafic and ultramafic cumulates beneath magmatic arcs may also help make arc bulk compositions more felsic [Kay and Kay, 1993; Jull and Kelemen, 2001]. The asthenosphere-lithosphere boundary beneath magmatic arcs probably lies close to the Moho, as suggested by the presence of low-velocity regions in the uppermost mantle (Figure 6a). Low-viscosity asthenospheric mantle circulating beneath the subarc Moho would provide an extremely favorable setting for dense mafic and ultramafic cumulates to sink (Figure 19). Sinking of mafic cumulates into convecting asthenosphere would also lead to a more felsic bulk composition for juvenile arcs than expected from consideration of primitive arc basaltic compositions. Ducea  estimates that this mechanism becomes important when the crust is thicker than 20–25 km; if so, this would mark an important transition in the evolution of any arc-trench complex.
6.3. Back Arc Region
 The back arc region lies behind the magmatic arc, and it can show a wide range of magmatic and tectonic styles, depending on strain class [Jarrard, 1986]. Low-strain arcs (strain classes 1 or 2) are associated with back arc extension, whereas arcs with high strain (strain classes 6 or 7) may be associated with back arc folding and thrusting. Intermediate strain classes may be associated with a back arc region showing little or no tectonic or magmatic activity.
 Active extension, rifting and seafloor spreading, characterize the back arc regions above several subduction zones: the Mariana Trough behind the Mariana arc, the Lau-Havre Trough behind the Tonga-Kermadec arc, the North Fiji Basin behind the Vanuatu (New Hebrides) arc, the Manus Basin NE of New Guinea, and the East Scotia Sea behind the South Sandwich arc are excellent examples of back arc basins with active seafloor spreading. Spreading rates approximate the range observed for mid-ocean ridges: from fast (16 cm yr−1 in the northern Lau Basin [Bevis et al., 1995]) to slow (4 cm yr−1 in the Mariana Trough [Bibee et al., 1980]). Back arc spreading systems are remarkably similar to mid-ocean ridges. Fast spreading back arc basins have inflated ridge morphologies frequently underlain by geophysically imageable axial magma chambers [Turner et al., 1999], while slow spreading back arc basins have axial rift morphology and occasionally expose mantle and lower crustal sections [Hawkins et al., 1990; Stern et al., 1996; Ohara et al., 2002]. Hydrothermal vent fields are common along back arc basin spreading ridges, although the composition of fluids, deposits, and biota is distinctive.
 Extensional back arcs may rift as well as spread. Rifting is observed where the extensional regime propagates along the strike of the arc system, such as the northern Mariana Trough [Martinez et al., 1995] and the Havre Trough [Fujiwara et al., 2001], as well as for back arc basins that are in the initial stages of development (Okinawa Trough SW of Japan [Fabbri and Fournier, 1999] and the Sumisu Rift in the Izu arc [Taylor et al., 1991]). How much of a back arc basin is underlain by “spread” crust as opposed to rifted crust is usually controversial, particularly for crust on the margins of the basin. Similar to the situation for rifting continents, rifted back arc basins can be amagmatic or volcanic. Back arc basins that form by seafloor spreading have crustal thicknesses that are indistinguishable from that of normal oceanic crust [Bibee et al., 1980]; rifted back arcs have crust that is intermediate in thickness between spread crust and the thickness of the associated, unrifted arc.
 Lavas erupted at back arc basins are commonly referred to as back arc basin basalts (BABB), but this term masks differences between melts associated with different evolutionary stages. BABB erupted from spreading ridges are dominated by pillowed basalts that are compositionally similar to MORB [Hawkins and Melchior, 1985] but contain more water (Figure 18) and a significant “subduction component” (Figure 14) [Gribble et al., 1998; Newman et al., 2000]. BABB from back arc rifts, however, may be compositionally similar to those erupted from the affected arc. Rift volcanism commonly yields a bimodal assemblage of basalts and felsic lavas. Mafic end-members may be similar to arc lavas (as is found in the northern Mariana Trough and eastern Manus Basin), or they can be indistinguishable from basalts produced by back arc spreading (as is found in the Sumisu Rift [Gribble et al., 1998; Hochstaedter et al., 1990]). These variations in melt chemistry provide important clues to mantle flow beneath a maturing back arc basin. Mantle may be sequentially melted, first, beneath the spreading ridge and, second, beneath the arc, and this may be responsible for much of the depletion observed in arc lavas [Woodhead et al., 1993]. Proximity of the arc to the back arc spreading axis can also enhance melting and magma supply [Martinez and Taylor, 2002].
 Crustal shortening and compression characterizes the back arc regions of high-strain convergent margins, leading to the development of a system of retroarc foreland basins behind the magmatic arc (Figure 20d). The best Cenozoic examples are found behind the present Andean arc and in western North America [Jordan, 1995]. These are fold-and-thrust belts, both of which formed in response to subduction of young, buoyant lithosphere. The Andean system is still active, whereas the North American system, which produced the Sevier-Laramide structures of Cretaceous-Paleogene age, was extinguished when the East Pacific Rise subducted beneath California in mid-Tertiary time.
Figure 20. Tectonic variability found behind magmatic arcs, using the Bolivian Andes and the Lau Basin as examples. Dark shading denotes lithosphere; underlying white is asthenosphere. (a) Tectonic setting of Lau Basin (modified after Zhao et al. ), emphasizing lithospheric structures. Dashed box shows region of detail shown in Figure 20b. (b) Tectonic setting of Andes between 20°S and 24°S (modified after Yuan et al. ) emphasizing crustal structure. Area of detail shown in Figure 20d lies partly within the open-ended dashed rectangle. (c) Cross section across the Lau Basin, emphasizing lithospheric structure, at the latitude investigated by Ocean Drilling Progdram Leg (∼20°S) (modified after Hawkins ). Extension rate is from GPS measurements of Bevis et al.  for velocities of Tonga from Australia; note that the highest rate is at the northern end of Lau Basin (∼17°S), and the rate decreases to the south. (d) Cross section across Andean back arc region, emphasizing crustal structure, about 20°S (modified after Gubbels et al. ). Shortening rate is from GPS measurements of Bevis et al. . Figures 20c and 20d have the same horizontal scale. There is no vertical exaggeration on any of the sections.
Download figure to PowerPoint