PII S0016-7037(01)00670-6

Tracing the evolving flux from the subducting plate in the Tonga-Kermadec arc systemusing boron in volcanic glass

PETER D. CLIFT,,1 ESTELLE F. ROSE,1 NOBUMICHI SHIMIZU ,1 GRAHAM D. LAYNE,1 AMY E. DRAUT,1 and MARCEL REGELOUS

2

1Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA2Max-Planck-Institut fur Chemie, Abteilung Geochemie, Postfach 3060, 55020 Mainz, Germany

(Received July24, 2000;accepted in revised form May4, 2001)

Abstract—The influence of fluid flux on petrogenesis in the Tonga-Kermadec Arc was investigated using ionmicroprobe measurements of B/Be and boron isotope ratios (11B/10B) to document the source and relativevolumes of the fluids released from the subducting oceanic plate. We analyzed young lavas from eightdifferent islands along the Tonga-Kermadec Arc, as well as glass shards in volcanic sediments from OceanDrilling Program (ODP) Site 840, which record the variations in the chemistry of Tonga magmatism since 7Ma. B/Be is variable (5.8–122), in young Tonga-Kermadec Arc lavas. In contrast, glass shards from;3 to4 Ma old volcanic sediments at Site 840 have the highest B/Be values yet reported for arc lavas (18–607).These values are too high to be related simply to a sediment influence on petrogenesis. Together with veryhigh d11B values (211.6 to137.5) for the same shards and lavas these data indicate that most of the B isderived from fluid escaped from the subducting altered Pacific oceanic crust, rather than from sediment. Highd11B values also reflect large degrees of isotopic fractionation in this cold fast subduction zone. Lowerd11Bvalues noted in the Kermadec Arc (17 to24.4) are related to the influence of sediment eroded from NewZealand and slower convergence. High fluid flux (B/Be) is synchronous in Tonga and the Marianas at 3 to 4Ma and may be related to acceleration of the Pacific Plate just prior to this time.

The timing of maximum B/Be at 3 to 4 Ma correlates with maximum light rare earth (LREE) and high fieldstrength element depletion. This suggests maximum degrees of partial melting at this time. Although thinningof the arc lithosphere during rifting to form the Lau Basin is expected to influence the arc geochemistry,variable aqueous fluid flux from the subducting plate alone appears capable of explaining boron and other traceelement systematics in the Tonga-Kermadec Arc with no indication of slab melting.Copyright © 2001Elsevier Science Ltd

1. INTRODUCTION

Understanding the processes that control arc petrogenesis isone of the outstanding problems in modern petrology, but adifficult task due to the large number of potential variablescompared to the relatively straight forward decompressionmelting of a mid ocean ridge system. In most subductionsystems the arc volcanic front is located over the dewateringslab at around 100 km depth (Tatsumi et al., 1983). The fluid andamphibole, generated by hydration of the mantle wedge above asubducting oceanic plate, is carried down and under the arc due toconvection in the wedge. As it does so the material is warmed andmelting begins to occur above about 1000°C (Tatsumi et al.,1983). Further melting occurs due to decompression during asthe-nospheric upwelling in a melt column under the arc volcanoes.

Petrogenetic models for subduction systems have focused ontwo key variables in controlling petrogenesis. In one set ofmodels, typified by Plank and Langmuir (1988), lithosphericprocesses dominate in constraining the extent of partial meltingby limiting the degree of upwelling of the mantle wedge abovethe subducting plate. In the other set of models the geochemicalcharacter of an arc’s output is mostly controlled by the natureof the fluid flux from the subducting plate. Within this lattergroup of models two approaches have been taken, those inwhich the fluid is variable and comprises of components from

the altered oceanic crust (AOC) and the sedimentary cover(Stolper and Newman, 1994), and those models in which thesedimentary cover dominates (Hole et al., 1984; Elliott et al.,1997; Plank and Langmuir, 1998).

Different methods have been used to demonstrate slab in-volvement in arc magmatism. The high concentrations of largeion lithophile elements (LILE) relative to high field strength(HFSE) and rare earth elements (REE) in arc lavas, comparedto midocean ridge basalts, show that the LILE budget of arclavas is typically derived from subducted materials (e.g.,Pearce, 1983). For example, oceanic basalts have uniformCe/Pb ratios of about 25 (Sun and McDonough, 1989), whereasmost arc lavas have much lower Ce/Pb, due to addition of Pbfrom subducted sediment and altered oceanic crust (e.g., Milleret al., 1994). Similarly, much of the Ba, U, Rb, Sr, as well asB, in arc lavas is also derived from the subducted slab.

Isotopic data also demonstrate slab involvement in arc mag-matism. The cosmogenic isotope10Be has been found in sev-eral arcs (e.g., Brown et al., 1982; Tera et al., 1986; Morris etal., 1990), implying rapid sediment subduction and melting,because this element is radioactive with a half-life of only1.5 m.y. Similarly Sr and Pb isotopes, especially the207Pb/204Pb ratio, has been used as strong evidence for sedimentinvolvement in several arcs (e.g., Aleutians and Alaska; Kay etal., 1978). However, the influence of the altered oceanic crust(AOC), as well as the sediments, has also been highlightedusing Pb isotopes (Miller et al., 1994) and the boron system(Spivack and Edmond, 1987; Seyfried et al., 1984).

*Author to whom correspondence should be addressed (pclift@whoi.edu).

Pergamon

Geochimica et Cosmochimica Acta, Vol. 65, No. 19, pp. 3347–3364, 2001Copyright © 2001 Elsevier Science LtdPrinted in the USA. All rights reserved

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3347

If fluid flux is the principal control on petrogenesis thenvariations in this flux should have a major effect on arc geo-chemistry. Hole et al. (1984), and more recently Plank andLangmuir (1993) and Johnson and Plank (1999), suggested thatthe REE characteristics of many arc lavas are derived from thesediments being subducted in the adjacent trench. In the Mari-anas a relative negative anomaly in Ce seen in the lavas fromAgrigan Island was compared with a negative Ce anomaly inthe sediments on the Pacific Plate adjacent to this section of thearc. Indeed, Hole et al. (1984) suggested that the light REE(LREE) enrichment known from the Northern Marianas mayreflect subduction of a large thickness of hotspot-derived clas-tics, rather than a long lived chemical anomaly in the mantle, asproposed by Lin et al. (1990). If hypotheses such as these arecorrect then understanding the nature and controls on flux fromthe subducting oceanic slab is crucial to understanding theextent to which subducted sediment is recycled back into themantle, and the process of new crust formation at convergentplate margins.

In this paper we examine a single key variable in this system,the fluid flux from the subducting slab within the classic intra-oceanic Tonga subduction system using the boron system. Weexploit the temporal dimension of arc volcanic activity avail-able through the volcaniclastic record of the forearc basin tochart the evolution in slab flux to the arc volcanic front in termsof its magnitude and origin, and examine how this has changedsince 7 Ma.

2. SETTING OF TONGA ARC

The volcanically active Tofua Arc and Lau Basin represent aclassic arc and backarc basin system (Karig, 1970; Hawkins,1974; Gill, 1976), which has developed as a result of subduc-tion of the Pacific lithosphere under the eastern edge of theIndian-Australian plate. Convergence between the two plates is;5.3 cm/yr. in the Kermadec Arc (DeMets et al., 1990),peaking at;20 cm/yr. in northern Tonga because of the addi-tional spreading in the Lau Basin (Pelletier et al., 1998). Sub-duction commenced in this region at approximately 45 Ma(Herzer and Exon, 1985) since which time the plate marginappears to have been in a constant state of subduction erosion(Clift and MacLeod, 1999). This means that sediment on thePacific Plate has not accreted during some periods and thenbeen eroded later to produce pulses in sediment flux to the arc,but has instead been always fully subducted, along with varyingamounts of the forearc basement and sedimentary cover. Thepresent arc system dates from the Late Miocene, when the LauBasin was generated by rifting of the Tonga Arc (Parson et al.,1992). Rifting had started by at least 7.0 Ma, although thegeneration of new backarc crust may postdate this by as much1.4 m.y. A minimum age of 5.6 Ma is given for the oldestigneous crust in the new basin at ODP Site 834 (Fig. 1;Shipboard Scientific Party, 1992). The original Miocene arc isnow preserved on the western side of the Lau Basin as aremnant arc called the Lau Ridge. Activity on the new “TofuaArc” dates from about 3.0 Ma (middle Pliocene; Tappin et al.,1994; Clift et al., 1995). The Tonga Platform has thus been ina forearc position to both the Miocene Tonga and the newTofua Arcs, and so in a position to record the volcanic productsof each. Volcanism in the Lau Basin is not thought to have

contributed to the sedimentation on the Tonga Platform, due tothe great difference in water depths and the submarine charac-ter of the volcanism in that area (Clift et al., 1995).

3. PREVIOUS WORK

The Tonga-Kermadec Arc system has been the target ofmany geochemical studies. Pearce et al. (1995) highlighted thesubduction input by showing that basalts from the Eastern LauSpreading Center have much higher enrichment in LILEs rel-ative to the Central Lau Spreading Center (Fig. 1). Traceelement and Th, Sr and Pb isotopic data from the modern arcvolcanic centers have been used to argue that in northern Tongamuch of the slab flux is from the altered oceanic crust, con-trasting with the more sediment-dominated signature seen inthe Kermadec Arc to the south, where a greater thickness of

Fig. 1. Bathymetric map of the Tonga-Kermadec Arc showing thelocation of the principle bathymetric features, scientific drill sites andactive volcanic centers described in this study. ELSC5 Eastern LauSpreading Center, CLSC5 Central Lau Spreading Center.

3348 Peter D. Clift et al.

sediment is being subducted (Regelous et al., 1997; Ewart etal., 1998). Clift and Vroon (1996) used Pb, Nd and Sr isotopicdata from the volcaniclastic record to support the idea of atemporary peak in the subducted sediment contribution to thearc following the start of backarc rifting. There was a delay of;3 m.y. between the onset of rifting and the peak of therecognized sediment pulse, a figure that agreed well with theconclusions of Regelous et al. (1997), Turner and Hawkes-worth (1997) and Turner et al. (1997) that 2 to 4 m.y. isrequired for Pb in the material subducted at the trench to beerupted at the adjacent volcanic front.

Clift and Dixon (1994) analyzed single glass grains fromvolcanic turbidites cored at ODP Site 840 for a series of major,trace and REEs (Fig. 1). The most primitive glass shardsshowed that HFSEs became increasingly depleted in the mostincompatible elements between 7.0 and 4.0 Ma. This trend wasreflected in the major element chemistry as a fall in the alkalielement contents, interpreted to reflect variations in the degreeof partial melting due to thinning of the arc lithosphere duringLau Basin rifting (cf., Plank and Langmuir, 1988).

4. NATURE OF THE VOLCANICLASTIC RECORD

The sedimentary sequences that accumulate in forearc basinsact as valuable records of the tectonic and magmatic evolutionof the active plate boundary. Chemical analysis of redepositedvolcaniclastic sediments has provided a new way of tracing arcchemical evolution (e.g., Hiscott and Gill, 1992) which can bereadily correlated with tectonic events through dating of inter-calated pelagic sediments. This approach hinges on the realiza-tion that the shards that comprise volcaniclastic sediments aresamples of the volcanic glass erupted at the arc volcanic frontat the time of deposition of the tephra or turbidite. Although notsingle-event tephra horizons, the turbidites incorporate shardsfrom a modest number of eruptions that occurred over geolog-ically short time spans. The redeposited character of the sedi-ments at ODP Site 840 was discussed in detail by Clift (1994).A summary log is provided here to demonstrate large-scalevariability appropriate to this study (Fig. 2A). The sands them-selves comprise a mixture of volcanic chemistries, dominatedby dacitic compositions, but with minor andesite, basaltic an-desite, plagioclase and pyroxene grains (Clift et al., 1995).Texturally the grains are fresh (Fig. 2B) and show a vesicularfabric, most strongly in the high silica grains.

Seismic data from Tonga (Tagudin and Scholl, 1994; Austinet al., 1989) suggest that large talus aprons form around the arcvolcanoes, while further towards the trench the sediment se-quences onlap the outer arc high from which substantialamounts of material may be episodically eroded (Clift et al.,1998). In practice sampling close to the arc centers, effectivelywithin the debris aprons, ensures that the sediments representthe local volcanic output, with minimal “contamination” fromalong-strike or from eroded older sequences along the outerforearc high. ODP Site 840 is located close enough to Ata (;25km) that it can be used as record of the activity of that center(Fig. 1).

Unfortunately even at ODP Site 840 the volcaniclastic recordof turbidites and debris flow conglomerates is not completebecause of a hiatus from;3.5 to 4.5 Ma (Fig. 2). Since riftingof the Lau Basin separated the original Miocene arc from the

forearc basin a volcaniclastic hiatus developed in the forearcstratigraphy (e.g., Clift et al., 1995). Fortunately this gap in therecord of the forearc can be filled by that preserved at ODP Site834, which represents the oldest submarine volcanism withinthe Lau Basin, effectively the arc volcanic front, at a time whenthere was no subaerial arc.

5. BORON IN SUBDUCTION ZONES

Boron has become an increasingly important element intracing the influence of recycling material in island arc systems.Studies of fluids vented from subduction accretionary prismsindicate that most boron is held in an exchangeable form in thesedimentary cover or altered oceanic crust (AOC), but not inthe mantle wedge (e.g., Spivack et al., 1987; Morris et al.,1990). Boron is lost mostly from the subducting plate at lowtemperatures (e.g., You et al., 1993, 1995), but a minor amountis structurally bound into minerals and can be transferred to themantle. A complete mass balance requires that some exchange-able boron is subducted to relatively deep levels, especially incolder, faster subduction zones (Bebout et al., 1999). If sub-duction is rapid then boron dissolved in sediment pore-watersmay be subducted further than the normal shallow regionsbelow forearcs.

Be is less fluid mobile than boron and is subducted to deeperlevels, behaving much like a LREE (Tatsumi and Isoyama,1988). Although partition coefficients of B and Be can differ bytwo orders of magnitude under certain conditions (Chaussidonand Libourel, 1993) the correlation of B/Be with other ratiosindicative of slab-derived fluids (e.g., Ba/Ce, Rb/La) demon-strates that these elements have very similar mineral-melt par-tition coefficients in most subduction zone environments(Ryan, 1989; Ryan and Langmuir, 1993). Partial melting andfractional crystallization processes do not therefore signifi-cantly fractionate B from Be, and variations in the B/Be ratio inarc lavas are controlled primarily by differences in the slabinput to the mantle sources of the lavas. Most arc lavas havesignificantly higher B/Be ratios than midocean ridge and oce-anic island basalts, because boron is added to the source of thearc lavas by fluids derived from the subducting slab. The B/Beratio is thus a useful indicator of the amount of slab-derivedboron in arc lavas (Morris et al., 1990; Edwards et al., 1993;Gill et al., 1993; Hochstaedter et al., 1996). Correlations of theB/Be ratio with10Be/9Be in some arc lavas (Morris et al., 1990;Leeman et al., 1994) suggest that boron is derived at least inpart from subducted sediment, and is rapidly transferred fromthe subducting slab to the surface in lavas (within about 5half-lives of 10Be, or 7.5 m.y.). Morris et al. (1990) note thatmodern oceanic sediments have B/Be ratios of 50 to 60, andoccasionally close to;100, so that values exceeding that levelrequire input from an additional boron reservoir, probably theAOC.

Boron isotopes have been used successfully to trace thesource of fluids from the subducting plate in a number of arcsystems (Ishikawa and Nakamura, 1994; Ishikawa and Tera,1999). Source resolution is possible because of the large boronisotopic differences between the surface reservoir for boron(i.e., especially between deep sea sediments, the AOC and themantle wedge). Although sedimentd11B is variable, mostlydepending on sediment composition (from –17.0‰ to14.5‰

3349Boron isotopic tracing of slab-flux in Tonga

for non-carbonate lithologies; Ishikawa and Nakamura, 1993),the mantle has relatively homogenous negative values ofd11B(210 6 2; Chaussidon and Marty, 1995). In contrast, AOCshows positive values ofd11B (between 0.1–9.26 0.4‰;

Spivack and Edmond, 1987; ranging up to124.9‰; Smith etal., 1995). These features make boron a valuable tracer of theorigin of slab-derived material (melt and/or fluid) in subductionzones.

Fig. 2. (A) Simplified stratigraphy at ODP Site 840 located on Tonga Forearc close to Ata. Note the expanded,coarse-grained record from;8 to 5 Ma and the more condensed record since 3.5 Ma, following the rift-related hiatus. (B)Backscattered electron-microprobe image of the glass shards, showing the vesicular nature of some of the materials. Darkcolors correspond to more mafic grains, although dacitic material dominates.

3350 Peter D. Clift et al.

5.1. Analytical Conditions

The Cameca ims 3f at Woods Hole Oceanographic Institu-tion was set-up for MRP51800, energy slit, field aperture andcontrast aperture settings following procedures developed byChaussidon et al. (1997). The primary beam is accelerated to 10keV, with a secondary accelerating voltage of 4.5 keV andproduces a beam size of 20 to 40mm on the sample. Back-ground interference levels are low,;0.05 ppm, and matrixeffects are negligible (Chaussidon et al., 1997). Replicate anal-yses of a silicate glass standard (GB4) have producedd11B 5212.76 0.8‰ (2s standard error, n5 9), consistent with theaccepted value ofd11B 5 12.80‰ (Chaussidon et al., 1997).The precision of individual analyses, involving 60 cycles ofmeasurements (counting times were 15 s for10B and 8 s for11Bfor each cycle), ranges from60.6 per mil to62.2 per mil (2s),comparable to those of Chaussidon et al. (1997).

Variations in boron isotope ratios are described by the no-tation d11B where

d11B5 F 11B/10B Sample11B/10B Standard

2 1G 3 1000 (1)

The standard used is NBS, the Searles Lake evaporite (Mc-Mullen et al., 1961).

Results of the ion probe boron analysis are shown in Tables1 and 2. In addition, subaerial volcanic rock specimens fromthe volcanic islands of the Tonga-Kermadec Arc were analyzedfor a range of REEs and trace elements using the Cameca ims3f microprobe. The results from these rocks are given in Table2, with analytical uncertainties in Table 3.

6. RESULTS

B/Be varies significantly along the strike of the arc, with arange of values seen at each volcano (Fig. 3). The Tonga andKermadec segments of the arc have similar B/Be ratios around15 to 95. B/Be never exceeds 122. Raoul in the Kermadecsshows the greatest range and highest values of measured B/Be(ranging 25–122), although Hunga Ha’apai and Fonualei(55–76 and 29–86 respectively) also appear to be rather highervalues than Tafahi, Tofua and Ata (14–52, 28–33, and 23–27respectively). Temporal variation in B/Be measured in thevolcaniclastic record (Fig. 4) shows that in fact modern volca-nic B/Be values are small (,130) compared to earlier activity(values up to 520 at 3–4 Ma). The record at ODP Site 840shows a well developed increase in B/Be values from about 70at 7 Ma, to about 300 at 5 Ma. B/Be reaches peak values of.400 at 3 to 4 Ma, following arc break-up. The highest, buthighly variable, values (.500) are noted at the base of thesection at ODP Site 834. Since 3 Ma there appears to be agradual decline in B/Be values to the present day value of 27 atAta, most notably at ODP Site 840. B/Be values measured inthe youngest volcaniclastic layers are consistent with the valuemeasured from modern Ata, the closest modern volcanic center(20–30). Globally modern Tonga-Kermadec has a relativelyhigh range of B/Be values (6–122), although similar to thoserecorded from Nicaragua (Reagan et al., 1994), Papua NewGuinea (Gill et al., 1993), and the Bismarck Arc (Morris et al.,1990).

Although there is a temporal variation in B/Be at ODP Site

840, the isotopic character of the volcanic glass does not showa systematic change with time through the drilled stratigraphy(Fig. 5). d11B values seen since 7 Ma at ODP Site 840 aremostly strongly positive (ranging from 0.6–37‰), with occa-sional negative excursions at 1.7 Ma and 3.6 Ma (21.3 and211.6, respectively).d11B values appear to be lower (fromabout11 to 115‰) at the base of the section and then rise tohigher values (112 to125‰) by;5.5 Ma. Since that time thevalues have been constantly positive and high, typically 10 to30. Some single intervals show significant spread of isotopicvalues, e.g.,d11B values from211.6 to132.8‰ in differentgrains in a single turbidite sand dated at 3.6 Ma. These analysesrepresent the first boron isotopic data from volcaniclastic de-posits published. Thed11B values are considerably more pos-itive than existing boron isotopic data from island arc lavas(usually26 to 17‰; Peacock and Hervig, 1999).

d11B is dominantly positive along length of the modern arc(ranging from244‰ to 30.8‰), but is highest in central andsouthern Tonga (from24.4 to130.8‰, Fig. 6A). Raoul inparticular has lowerd11B, but with a spread of values from24.4 to 117.1‰. There is a good deal of scatter at each arcvolcano, even within analyses conducted on the same sample.This trend may reflect either small scale magmatic variabilityor limited alteration. Rain water hasd11B values of13.8 to19.1‰ (Rose, 1999), however its effectiveness as a contami-nating agent in the case of the subaerial lavas is very muchrestricted by the B concentration of 1 to 5 ppb.

7. DISCUSSION

7.1. Diagenetic Alteration

Since boron is a very water-mobile element there is clearlyconcern regarding the meaning of boron analyses made involcanic glass shards as small as 40mm across, which havespent significant amounts of time surrounded by seawater anddiagenetic fluids. This concern is especially strong for silicicgrains which tend to be more vesicular and thus have a highsurface area/volume ratio. Several factors now suggest that thealteration of the glass grains considered here is low, at least inthe upper part of the drilled section at ODP Site 840, and thatthe boron is representative of the magma source and not burialdiagenesis.

Visual inspection of the grains reveals clear pristine grains,unclouded by hydration, except in rare cases towards the baseof ODP Site 840. However, chemical alteration of mobileelements may begin before visible change has occurred, andthis may be seen as low analytical totals in electron probeanalysis. In arc volcanic rocks low analytical totals can arisefrom an indigenous volatile content in the melt or from subse-quent hydration. Sobolev and Chaussidon (1996) estimate thatprimary subduction melts contain 1 to 3% H2O, values that maybe expected to rise during crystal fractionation (e.g., Burnhamand Jahns, 1962), although much of this may be lost by degas-sing during eruption. In a varied suite of lavas evolving underdifferent conditions of volatile degassing, and perhaps differingprimary water contents, a range of volatile contents might beexpected over much of the compositional spread, with a broadcorrelation between volatile-loss measured by [100%—Total],and silica content. Figure 7A shows this to be the case. The datahas a sloping upper bound which corresponds to the effective

3351Boron isotopic tracing of slab-flux in Tonga

Table 1. B and Be concentration and isotopic data from SIMS analysis of single tephra particles from ODP Sites 834 and 840.

Age (Ma) Be (ppm) B (ppm) %2s 11B/10B d11Boron Volatile %

840A–1–1, 102cm 0.35 0.198 4.70 0.39 3.911 9.097 2.49840A–1–1, 102cm 0.35 0.198 4.70 0.78 3.906 6.814 6.76840A–1–1, 102cm 0.35 0.198 4.70 — — — 8.30836–2–6, 38cm 0.40 0.012 2.31 — — — 3.05840B–1–CC, 33cm 0.40 0.020 2.59 — — — 3.37840A–1–2, 75cm 0.43 0.018 2.17 — — — 0.10840A–1–2, 75cm 0.43 0.018 2.17 0.47 3.963 22.497 15.34840A–1–2, 75cm 0.43 0.018 2.17 0.70 3.936 14.490 0.40840A–1–2, 75cm 0.43 0.018 2.17 — 11.17840A–1–3, 45cm 0.45 0.021 2.57 0.73 3.979 26.634 0.22840A–1–3, 45cm 0.45 0.021 2.57 — — — 2.95840A–1–3, 45cm 0.45 0.021 2.57 — — — 1.71840B–1–CC, 33cm 0.45 0.020 2.59 — — — 0.86840B–1–CC, 33cm 0.45 0.020 2.59 0.41 3.937 15.742 2.66839–2–1, 44cm 0.50 0.025 1.19 — — — 7.90836–3–6, 123cm 0.60 0.019 2.19 — — — 4.39836–3–6, 123cm 0.60 0.017 1.31 — — — 5.75838–3–2, 147cm 0.75 0.025 0.79 — — — 6.48839–3–4, 22cm 0.75 0.026 0.97 — — — 5.34835–4–2, 121cm 0.80 0.025 1.18 — — — 7.16838A–5–5, 85cm 0.80 0.022 1.25 — — — 5.96839–4–4, 35cm 0.95 0.028 1.64 — — — 6.80834–3–1, 72cm 1.00 0.020 3.69 — — — 2.89837–4–6, 67cm 1.00 0.017 3.72 — — — 3.76835–8–6, 25cm 1.45 0.024 2.98 — — — 5.21837–6–4, 132cm 1.60 0.023 0.77 — — — 2.33839–10–3, 60cm 1.68 0.016 3.69 — — — 5.65839–11–6, 82cm 1.70 0.023 0.59 — — — 3.95835–14–6, 84cm 1.70 0.037 5.04 — — — 4.79837–7–4, 33cm 1.70 0.022 1.30 — — — 3.10838–7–2, 89cm 1.70 0.016 0.62 — — — 4.82838–7–2, 89cm 1.70 0.024 1.96 — — — 3.76840B–3X–CC, 10 cm 1.70 — — 0.83 3.946 5.869 4.93840B–3X–CC, 10 cm 1.70 — — 0.88 3.928 1.468 7.50840C–1H–2, 29cm 1.70 0.013 3.14 0.33 3.966 23.210 4.93840C–1H–2, 29cm 1.70 0.013 3.14 — — — 7.50840C–1H–2, 114cm 1.73 — — 0.66 3.917 21.329 3.17839–15–1, 79cm 1.73 0.014 0.91 — — — 3.01835–16–3, 100cm 1.73 0.028 1.40 — — — 6.33839–18–CC, 6cm 1.74 0.017 0.98 — — — 3.67839–21–CC, 33cm 1.74 0.030 0.88 — — — 3.82837–8–2, 49cm 1.90 0.018 0.73 — — — 1.83838–11–5, 30cm 1.95 0.017 5.32 — — — 5.39837–9–1, 53cm 2.00 0.015 4.01 — — — 4.24837–9–3, 93cm 2.05 0.017 1.19 — — — 2.55840C–2–1, 37cm 2.20 0.012 3.87 — — — 5.22840C–2–1, 37cm 2.20 0.012 3.87 0.29 3.990 29.439 5.65840C–2–1, 37cm 2.20 0.012 3.87 0.37 3.991 15.940 5.76840C–3–3, 2cm 2.70 0.011 4.79 0.55 3.960 21.650 4.18840C–3–3, 2cm 2.70 0.011 4.79 — — — 2.31840C–3–3, 2cm 2.70 0.011 4.79 — — — 4.92834–6–2, 35cm 2.90 0.021 3.15 — — — 6.28840C–4–6, 34cm 3.00 — — 0.61 3.967 23.541 1.09835–15–6, 83cm 3.05 0.035 5.59 — — — 7.24834–7–4, 67cm 3.30 0.033 3.44 — — — 3.09840B–10X–CC, 4cm 3.40 0.013 5.48 0.23 3.967 23.541 6.29834–9–3, 98cm 3.50 0.029 7.27 0.22 3.922 12.019 7.86834–9–3, 98cm 3.50 0.029 7.27 — — — 6.74834–9–4, 95cm 3.52 0.018 4.91 — — — 6.16834–9–4, 95cm 3.53 0.018 4.91 0.36 3.988 28.863 9.50834–9–4, 95cm 3.53 0.018 4.91 0.76 3.966 23.163 7.83834–9–6, 92cm 3.53 0.018 8.75 0.45 3.976 18.292 7.58834–9–6, 92cm 3.55 0.018 8.75 0.63 3.983 20.046 7.58834–10–3, 102cm 3.57 0.022 3.73 0.55 4.003 32.853 5.19834–10–3, 102cm 3.57 0.022 3.73 0.53 3.882 211.633 6.90834–10–3, 102cm 3.57 0.022 3.73 — — — 7.84834–10–CC, 13cm 3.58 0.016 3.39 — — — 4.10

(Continued)

3352 Peter D. Clift et al.

upper limit of SiO2 in a granitic liquid when subject to avariable addition of volatiles. This plot allows very anomalouslow totals at a particular silica content to be identified. This plotdoes not sort bad from good analyses but instead only identifiesthose grains that are almost certainly significantly hydratedduring diagenesis. Although diagenesis may affect grainsthroughout the section, the temporal distribution of those withanomalously low totals shows a clear pattern of strongestalteration of the oldest grains, as might be expected (Fig. 7B).There is no correlation between high B/Be and the zone at thebase of the section at ODP Site 840 where alteration is mostcommon. This pattern is consistent with, but not conclusive of,a primary origin for the boron in the upper part of the section.

There is no apparent relationship between total volatile con-tent and boron isotopic ratios (Fig. 7C). However, there is somecorrelation between lowerd11B and higher alteration at thebase of ODP Site 840 (Fig. 7D), consistent with the observationthat diagenesis and alteration of volcanic glass tends to de-creased11B values (Ishikawa and Nakamura, 1993). In Figure7D the amount of excess volatiles beyond the maximum thatcan reasonably be associated with fractional crystallization is

compared withd11B. The least altered grains (i.e., with highnegative excess volatiles,,25%) show a range ofd11B valuesbetween 20 and 29‰. However, the lowestd11B values (,6‰)are found only in grains with the higher volatile excess (.24%excess volatile), i.e., the most altered grains tend to have lowerd11B values. The higher values that dominate the upper part ofthe hole can be considered to be the most magmatic.

Volcaniclasticd11B values in sediments younger than 5.5Ma, typically with the lowest excess volatile content, are closeto thosed11B values measured from the subaerial volcanicrocks of adjacent Ata (8.6 and 16.4‰), consistent with theircomposition being close to magmatic values. If thed11B involcanic sediments younger than;5.5 Ma had been affected bydiagenesis then these might be expected to have fractionated tod11B values more negative than those we can be reasonablycertain are primary and magmatic. We conclude that for sedi-ments younger than 5.5 Mad11B may be considered to be closeto original magmatic values.

It is noteworthy that the time of highest B/Be (3–4 Ma)follows backarc rifting and is coincident with a similar peak inthe Marianas (Clift and Lee, 1998). This peak may be a re-

Table 1. (Continued)

Age (Ma) Be (ppm) B (ppm) %2s 11B/10B d11Boron Volatile %

834-10-CC, 13cm 3.58 0.016 3.39 0.38 3.913 9.535 7.66834–10–CC, 13cm 3.58 0.016 3.39 0.41 3.932 14.420 6.24834–1o–CC, 13cm 3.58 0.016 3.39 — — — 3.65834–11–1, 40cm 3.70 0.018 10.84 — — — 6.00834–12–2, 88cm 3.80 0.014 7.36 0.39 3.962 22.275 9.29834–12–2, 88cm 3.80 0.014 7.36 6.52834–12–2, 88cm 3.80 0.014 7.36 — — — 6.22840B–11–1, 122cm 4.50 0.020 4.64 0.23 3.999 31.802 8.10840B–12–5, 49cm 5.00 0.022 2.08 1.00 3.964 9.157 3.92840B–12–5, 49cm 5.00 0.023 0.41 — — — 2.81840B–12–5, 50cm 5.00 0.026 2.09 — — — 2.98840B–13–1, 137cm 5.05 0.023 4.04 — — — 6.48840B–13–3, 137cm 5.05 0.023 4.04 0.29 3.989 29.101 3.70840B–13–3, 137cm 5.05 0.023 4.04 6.48840B–13–3, 137cm 5.05 0.023 4.04 3.70840C–5–CC, 7cm 5.08 0.022 3.39 0.60 3.957 20.800 4.53840C–5–CC, 7cm 5.08 0.022 3.39 0.58 4.022 37.576 6.71840C–6–4, 123cm 5.09 0.025 4.20 — — — 4.74840C–7–4, 128cm 5.11 0.018 4.59 — — — 2.71840B–17X–CC, 2cm 5.13 0.034 4.74 — — — 2.71840B–17X–CC, 2cm 5.13 0.034 4.74 — — — 0.15840C–8–1, 97cm 5.13 0.020 4.33 — — — 0.00840C–9–1, 2cm 5.25 0.013 2.37 0.38 3.991 29.758 3.10840C–10–1, 3cm 5.30 0.014 2.39 — — — 3.37840C–12–2, 132cm 5.33 0.019 3.13 — — — 6.56840B–26–1, 60cm 5.37 0.019 3.80 — — — 6.49840C–13–5, 79cm 5.38 0.014 3.15 — — — 6.49840B–28–1, 77cm 5.38 0.018 6.35 0.41 4.026 24.877 6.17840C–13–5, 79cm 5.38 0.014 3.15 0.59 3.973 17.485 5.19840B–30–1, 12cm 5.40 0.013 1.85 0.59 3.923 12.192 2.96840B–33–1, 21cm 5.65 0.019 2.13 0.88 3.905 7.475 5.89840B–33–1, 21cm 5.65 0.034 1.46 — — — 5.09840B–34X–1, 24cm 5.80 0.025 0.64 — — — 5.80840B–37X–1, 13cm 5.90 0.043 1.62 — — — 3.22840B–39–1, 53cm 5.98 0.020 2.50 — — — 12.10840B–39–1, 53cm 5.98 0.023 2.89 — — — 2.06840B–46X–2, 9cm 6.30 0.041 3.48 0.71 3.960 8.125 0.62840B–46X–2, 9cm 6.30 0.030 2.03 — — — 8.39840B–52X–2, 5cm 6.53 0.023 1.89 0.32 3.942 17.097 5.51840B–52X–2, 5cm 6.53 0.023 1.89 0.32 3.930 0.604 5.98840B–52X–2, 5cm 6.53 0.035 0.91 0.80 3.917 3.262 4.71

3353Boron isotopic tracing of slab-flux in Tonga

Tab

le2.

Tra

ceel

emen

t,R

EE

and

Ban

dB

eco

ncen

trat

ion

and

isot

opic

data

from

SIM

San

alys

isof

volc

anic

rock

sfr

omth

eT

onga

-Ker

mad

ecA

rc.

Isla

ndS

ampl

e#

LaC

eN

dS

mE

uD

yE

rY

bLi

Be

BT

iR

bS

rY

Zr

Nb

Ba

B/B

e%

2s11 B

/10 B

d11 B

Taf

ahi

T11

40.

441.

051.

110.

540.

531.

340.

901.

211.

200.

1642

8.64

939

4.4

0.8

139.

01.

73.

50.

215

21.7

52.6

80.

423.

9325

14.6

1T

afah

iT

114

0.73

1.52

1.05

0.39

0.54

0.76

0.56

0.73

0.25

0.12

582.

183

341.

00.

912

7.1

2.3

4.6

0.13

714

.117

.35

0.52

3.95

3819

.10

Taf

ahi

T11

40.

380.

650.

700.

380.

330.

770.

620.

741.

070.

1114

1.64

461

8.4

0.7

103.

72.

03.

70.

117

11.7

14.7

60.

713.

9554

13.0

5F

onua

lei

Fon

–31

2.78

6.49

4.14

1.35

2.08

2.02

1.19

1.72

6.71

0.10

486.

725

1221

.66.

518

7.1

6.6

21.4

0.37

214

8.8

64.1

70.

643.

9479

17.6

0F

onua

lei

Fon

–31

2.38

5.55

3.29

1.06

1.44

1.45

0.94

1.05

3.99

0.10

734.

126

1069

.63.

815

9.7

8.3

17.2

0.24

110

8.6

38.4

70.

683.

9497

11.6

1F

onua

lei

Fon

–31

2.87

6.61

4.35

1.23

1.18

1.82

1.53

1.57

3.15

0.07

265.

014

6.5

76.6

5.1

13.6

0.36

554

.469

.08

0.83

3.96

7216

.07

Fon

uale

iF

on–3

12.

115.

634.

311.

432.

791.

971.

381.

706.

920.

0844

6.84

112

04.0

6.4

90.8

12.8

33.1

0.56

717

2.8

81.0

2F

onua

lei

Fon

–8–6

93.

087.

604.

681.

562.

821.

971.

241.

3911

.10

0.09

228.

000

864.

58.

691

.116

.441

.40.

671

230.

586

.81

0.39

3.93

3614

.89

Fon

uale

iF

on–8

–69

3.17

0.01

610.

481

850.

54.

425

.512

.54.

70.

033

11.1

29.8

30.

363.

9382

16.0

6F

onua

lei

Fon

–8–6

90.

881.

871.

190.

280.

610.

350.

220.

188.

990.

0958

3.27

04.

534

.53.

29.

30.

795

27.0

34.1

10.

363.

9427

17.2

3T

ofua

Tof

ua–3

21.

195.

434.

211.

601.

592.

321.

421.

711.

900.

0871

2.51

625

19.5

3.1

141.

78.

219

.60.

202

89.8

28.8

70.

263.

9952

30.7

7T

ofua

Tof

ua–3

22.

005.

414.

381.

650.

832.

251.

531.

621.

750.

1134

3.83

518

97.6

3.8

121.

511

.921

.90.

176

93.0

33.8

10.

823.

9404

16.6

4T

ofua

Tof

ua–3

21.

243.

212.

230.

690.

941.

241.

020.

950.

483.

9726

23.9

6H

.H

a’ap

aiH

HT

op0.

962.

682.

651.

411.

012.

671.

922.

120.

160.

0794

4.34

223

9.1

0.9

129.

01.

31.

10.

026

28.5

54.6

70.

513.

9657

23.1

6H

.H

a’ap

aiH

HT

op3.

157.

815.

201.

641.

772.

691.

911.

950.

290.

0425

3.24

915

720

1.7

46.6

2.4

10.3

0.38

521

.476

.44

0.72

3.98

4327

.97

H.

Ha’

apai

HH

Top

0.50

1.26

0.92

0.57

0.49

0.72

0.55

0.78

0.04

0.03

591.

983

0.9

45.0

1.2

0.8

0.11

711

.555

.23

0.71

3.96

9223

.08

H.

Ha’

apai

HH

Top

0.51

3.96

4210

.60

H.

Ha’

apai

HH

Top

0.69

3.96

1014

.49

Ata

582–

8–4

3.91

10.3

76.

522.

192.

943.

001.

782.

120.

870.

1325

3.04

7523

.13

0.23

3.93

9616

.43

Ata

582–

8–4

2.58

6.52

4.67

1.80

1.44

2.53

1.59

1.96

4.62

0.11

543.

1233

27.0

60.

123.

9380

8.61

Ata

582–

8–4

3.97

10.9

79.

093.

102.

616.

014.

194.

83A

ta58

2–8–

43.

899.

365.

751.

381.

542.

001.

501.

44S

eam

ount

—2.

606.

815.

151.

732.

273.

062.

112.

464.

060.

0578

3.18

32.

850

.16.

012

.30.

300

48.6

55.0

40.

313.

9865

28.5

3S

eam

ount

—2.

156.

204.

171.

271.

841.

781.

291.

4910

.39

1.40

698.

166

16.3

90.0

7.9

27.3

8.12

810

5.4

5.80

0.42

3.97

2924

.03

Sea

mou

nt—

3.94

9.50

6.05

1.73

1.48

2.72

2.10

2.28

8.51

0.10

273.

938

7.9

73.9

9.1

16.7

1.16

074

.738

.36

0.45

3.91

8510

.99

Rao

ul23

374

1.12

3.62

3.07

1.38

1.08

2.83

1.86

2.16

2.99

0.10

0012

.215

5421

.35.

718

3.5

79.4

73.4

0.54

113

9.3

122.

160.

233.

9422

17.1

1R

aoul

2337

41.

504.

984.

792.

141.

634.

112.

522.

935.

210.

0849

8.15

127

87.2

5.3

109.

517

.249

.60.

404

130.

996

.02

0.47

3.94

2716

.25

Rao

ul23

374

1.06

3.52

3.65

1.40

0.84

2.93

2.22

2.29

2.71

0.09

718.

214

2838

.34.

180

.817

.436

.30.

254

69.9

84.6

13.

9494

6.83

Rao

ul23

374

2.18

6.64

5.12

2.02

1.98

3.45

2.22

2.62

4.73

0.10

182.

562

1282

.42.

712

4.4

6.2

14.7

0.11

141

.525

.17

0.46

3.88

492.

33R

aoul

2338

31.

795.

013.

531.

200.

781.

591.

221.

325.

260.

1274

3.72

913

99.8

2.7

132.

17.

020

.30.

241

49.4

29.2

73.

8857

2.54

Rao

ul23

383

3.89

834.

80R

aoul

2338

33.

9055

24.

36L’

Esp

eran

ce14

837

4.69

11.2

46.

882.

502.

063.

822.

362.

520.

030.

1458

8.02

113

94.3

1.8

155.

311

.968

.80.

215

171.

055

.00

0.45

3.92

0411

.48

L’E

sper

ance

1483

71.

154.

502.

641.

200.

692.

301.

822.

420.

060.

1574

3.41

860

3.3

1.1

147.

74.

65.

80.

039

40.9

21.7

2L’

Esp

eran

ce14

837

2.50

5.30

3.23

1.18

2.47

1.45

1.09

0.88

3354 Peter D. Clift et al.

gional feature of active margin volcanism at this time, or maybe linked to rifting of the respective backarc basins. If the 3 to4 Ma peak B/Be seen at ODP Site 840 were related to alter-ation, it would be remarkable that the same effect is seen notonly at ODP Site 834, but also in a different arc and differentborehole at different burial depths, but at the same stage in thearc’s evolution. We therefore suggest that the B/Be record ofthe tephra is not significantly damaged by the diagenesis, at

least back to;5.5 Ma. The super-high B/Be values noted (upto 607) must therefore represent a phase of extreme slab fluxinto the arc, which does not appear to have an equivalent in themodern oceans.

Table 3. (Analyses from 135–840A–1H–3, 45cm:Hf from 135–840B–35X–3, 22 cm)

ElementTypical % errorin count rates

Concentration(ppm)

Rb 2.4 7.040Sr 0.4 214.0Y 1.0 14.40Zr 0.9 60.00Nb 4.7 0.576Ba 1.0 188.0La 3.3 3.920Ce 2.3 9.780Nd 7.5 6.100Sm 12.1 1.670Eu 8.6 0.112Gd 9.3 2.900Hf 8.5 1.990Tb 9.3 0.406Dy 10.0 1.800Ho 9.1 0.522Er 11.5 1.550Tm 13.7 0.245Yb 11.5 1.560Lu 15.8 0.181Pb 2.6 1.160Th 26.7 0.174U 24.3 0.226

Fig. 3. Along-strike variations in B/Be in the Tonga-Kermadec Arc measured by SIMS analysis of volcanic glass and finegrained groundmass of modern arc volcanic rocks.

Fig. 4. Temporal variation in B/Be in the Tonga system as sampledat ODP Site 840. The record shows peak values at 3 to 4 Ma, just afterbreak-up of the Lau Ridge Arc. Values at the volcanic front havedecreased since 3 Ma to the modern values.

3355Boron isotopic tracing of slab-flux in Tonga

7.2. Volume of Fluid Flux

The high B/Be values in Tonga-Kermadec lavas are inter-preted to represent generally high volumes of modern slabfluid-flux, with up to 95% of the total boron being slab-derivedfluids. In contrast, B/Be values for the depleted upper mantleare,5 (Hochstaedter et al., 1996). The relative volume of slabfluid-flux charted by B/Be at ODP Site 840 shows long termvariability. The very high B/Be values of the sediments depos-ited at 3 to 4 Ma are about three times higher than reported forany other arc lavas yet analyzed, suggesting temporary dra-matic slab fluid-flux at that time (B from the slab comprising99% of total). They are also far in excess of known sedimentvalues which are typically 50 to 60, and reach a maximum of

;100 (Morris et al., 1990), requiring an additional source forthe excess boron. In contrast, known B/Be values from AOCrange from 100 to 400 (Fig. 7; Thompson and Melson, 1970;Seyfried et al., 1984).

B/Be values in volcanic glass are low in slow subductionzones, e.g., Aeolian, Campanian and Aegean Arcs (Morris etal., 1993, Clift and Blusztajn, 1999), or where the subductingplate is very young and hot, e.g., some parts of central America(Leeman et al., 1994) and Mexico (Hochstaedter et al., 1996).In contrast, values of B/Be are high in rapid subduction zones,or where the subducting plate is very old and cold, e.g., Mari-anas (Clift and Lee, 1998), Bismarck Arc (Morris et al., 1990).Although B/Be values are affected by several factors it appearsthat there is a first order correlation between B/Be and thethermal structure of the subduction zone, a consequence of thetemperature sensitivity of boron in subducting sediments. Ex-perimental data indicate that at elevated temperatures (200–350°C) most boron in subducting sediment has been mobilized(You et al., 1995), while more metamorphosed sedimentsshowing progressive loss, so that by;800 to 900°C the vastmajority of exchangeable boron in sediments or altered volca-nic rocks has been lost (Moran et al., 1992). By the time theslab reaches its melting point little excess boron remains. Inthese settings boron is subducted to greater depths, where itmay be introduced into the arc lava source. The Tonga-Ker-madec arc is likely to have a low thermal gradient because it issubducting relatively old and cold oceanic lithosphere andbecause the rate of convergence, especially in the north, is veryfast (;20 cm/yr.; Pelletier et al., 1998). The high B/Be in themodern Tonga-Kermadec arc, as well as in the volcaniclasticrecord, argues strongly for slab dewatering and against slabmelting under the arc volcanic front.

Fig. 5. Temporal variation in boron isotope ratios in the Tongasystem, showing a generally constant, positived11B ratio since 7 Ma.

Fig. 6. Along-strike variations in (A) boron isotope ratios in theTonga-Kermadec Arc, as measured by SIMS analysis of volcanic glassand fine grained groundmass of modern arc volcanic rocks. Variationsin (B) 230Th/232Th and (C)87Sr/86Sr for the Tonga Arc. Data fromRegelous et al. (1997) and Ewart et al. (1998).

3356 Peter D. Clift et al.

7.3. Fluid Sources

The isotopic boron compositions and B/Be values in volca-nic glass can be used as a guide to the source of the slab flux,because the mantle wedge, subducted sediments and AOC havevery different boron concentrations and isotope compositions.These signals are not transmitted directly to the surface becauseboron isotopic fractionation at low temperature produces fluidsderived from the slab with higherd11B than their source (e.g.,Peacock and Hervig, 1997, 1999). Those fluids could triggerthe melting of the mantle wedge and result in highd11B in arclavas. The degree of isotopic fractionation of the fluid expelledfrom the slab is a function of the temperature (Oi et al., 1989),

with the greatest fractionation occurring at low temperatures.The very highd11B seen along-strike is consistent with the coldstate of the Tonga-Kermadec subduction zone. Isotopic frac-tionation does not occur when melting occurs because there isno difference in compatible in siliceous melts. The very highd11B seen in this study are higher than any possible source andrequire transfer to the mantle wedge in aqueous solution, not bydirect melting of the subducting material.

Measuredd11B is rather higher than other known oceanic arcvolcanic compositions (e.g., Izu,d11B 5 11.2 to 17.3‰;Ishikawa and Nakamura, 1994; Halmahera,d11B 5 22 to14‰; Palmer, 1991; Martinique,d11B 5 26 to 12‰; Smith

Fig. 7. Diagram showing relationship betweend11B (origin of slab flux) and B/Be (volume of slab flux). High B/Be isalways associated with positived11B and a dominance of water derived from the altered oceanic crust. Black dots representanalyses from ODP Site 840 volcaniclastic sediment, circles represent modern arc lavas. Black squares representend-member compositions used in mixing calculations. Data from Tables 1 and 2. Mantle data from Ishikawa and Nakamura(1994) and Spivack and Edmond (1987), pelagic sediment data from Ishikawa and Nakamura (1993), altered oceanic crustdata from Thompson and Melson (1970) and Seyfried et al. (1984).

3357Boron isotopic tracing of slab-flux in Tonga

et al., 1997; Kurile,d11B 5 24 to 16‰; Ishikawa and Tera,1997). However, at low temperature (25°C) Peacock andHervig (1999) have demonstrated that fluids withd11B as highas120‰ can be expelled from sediments with initiald11B of0‰, even more positive than the115‰ d11B values measuredin natural samples by Spivack et al. (1987) for the exchange-able boron released from sediments in accretionary complexes.The boron isotopic composition of AOC, with initiald11Btypically between25 and110‰, can also undergo isotopicfractionation during dehydration. The fluid expelled will havemore positived11B value than the subducting slab (following aRayleigh distillation model, e.g., Peacock and Hervig, 1999).Although temperatures under the arc volcanic front might beexpected to be higher than those that yield the most extremeisotopic fractionation, highd11B values can be incorporatedinto arc volcanic rocks if the base of the hydrated mantle wedgeis dragged down with the subducting slab, as suggested byTatsumi (1989). This is more likely in faster, colder subductionsystems. Convergence at Tonga-Kermadec (20 cm/yr.) is sig-nificantly faster than Izu (8–10 cm/yr.), despite the similarthermal age of the oceanic lithosphere in each case, making thisarc most likely candidate for the most positived11B worldwide.

Positive d11B are reported in carbonates (Hemming andHanson, 1992; Gaillardet and Alle`gre, 1995), but their absencefrom the stratigraphy of the Pacific abyssal seafloor offshoreTonga (Burns et al., 1973) makes carbonates an unlikely can-didate for the source rock of the lavas with highd11B values.The pattern of positived11B along the modern Tonga arc mostlikely reflects a fluid-flux dominated by the dehydration of theAOC similar to what has been proposed for the Mariana Arc(Ishikawa and Nakamura, 1994, 1999), but with a greatercontribution of sediment dehydration in the southern extremewhered11B is lower. Li isotope work in the Izu Arc supportsthe hypothesis of fluid-flux from the AOC as being dominant inthat oceanic arc (Moriguti and Nakamura, 1998). Unfortunatelywe do not have boron isotopic data from the abyssal seafloorimmediately east of the trench. Instead we refer to the pelagicsediment data of Ishikawa and Nakamura (1993), made in thecentral Pacific. The wind-blown origin of deep sea clays makesthem particular uniform in composition over wide areas, so thatno significant error is expected as a result of this approxima-tion. Modern marine sedimentd11B values range from26.6‰(clay) to 110.5‰ (carbonate), although pelagic clays haveuniformly negative d11B values (Ishikawa and Nakamura,1993). The high positived11B values of the arc lavas contrastwith the low d11B values of the clays and argues against thembeing a major contributor to the boron budget of the arc.Without carbonate sediments the only appropriate boron reser-voir is the fluid expelled during the dehydration of the AOC.

Input of continental material eroded from New Zealandforms a plausible explanation for the more negatived11Bvalues measured in the Kermadec Arc. The Kermadec Trenchis largely filled with sediment, especially at its southern end, incontrast to the relatively empty Tonga Trench, in which eventhe igneous basement to the subducting plate is exposed inflexural fault scarps (Hawkins et al., 1999). In addition, becauseof the slower extension in the Kermadec backarc (HavreTrough) compared to Tonga the subduction zone might beexpected to have a warmer thermal gradient, so that the boron

isotopes are less susceptible to isotopic fractionation. Togetherthese factors drive the Kermadec lavas to lowerd11B values.

7.4. Mixing End-Member Compositions

Figure 6 shows the relationship betweend11B and B/Be.Although there is a lot of scatter at low values of B/Be, it isclear that all samples with high B/Be values also have highpositive d11B, and that all samples with negatived11B haverelatively low B/Be (,200). This trend is not due to seawatercontamination. A simple mixing model between seawater (B54.5 ppm,d11B 5 40‰, and B/Be5 16200) and an arbitraryend-member representing a melt produced under a low slabfluid-flux with B 5 1 ppm, d11B 5 0‰ and B/Be5 25,suggests that a contamination of 0.4% of seawater would in-crease drastically the B/Be (from 25–310), covering much ofthe measured B/Be variation. However, it would only increasethed11B value from 0 to10.7‰; this hardly represents 2% ofthe total measuredd11B variation. Similarly, the boron concen-tration would increase by less than 2%. The trend on Figure 6has a much steeper slope than the slope that seawater contam-ination would produce.

In contrast, if we mix the arbitrary end-member with AOC(B 5 13 ppm,d11B 5 18‰, Spivack and Edmond, 1987, andB/Be 5 350), it appears that a fraction between 1 and 45% ofAOC can explain most of the B isotopic and the B/Be varia-tions. Nonetheless, mixing the AOC directly would produced11B values slightly lighter than the ones measured in thetephra (Fig. 6). If AOC is dehydrated at the shallow levels ofthe subduction zone where the temperature is low, the fluidexpelled will have a heavierd11B than the initial AOC due toisotope fractionation (Palmer et al., 1987; Oi et al., 1989).Depending on the extent of dehydration and the relative parti-tion coefficient of B and Be between AOC and fluid phase, itpossible to produce the measuredd11B values (Fig. 6). Al-though we can not totally rule out the contribution of fluidsexpelled by dewatering of sediments, which could also producea fluid with higher d11B than the initial sediments, the Bisotopic fractionation required to reach the highd11B observedwould have to be more extreme. This implies that to satisfy theboron data from the tephra, the model requires preferentialinvolvement of the altered fraction of the subducting AOC,through derivation of a fluid.

In summary, the relatively highd11B of Tonga-Kermadeclavas, compared to other arc lavas, results from the high rate ofplate convergence and the Cretaceous age of the oceanic litho-sphere that result in a cold subduction zone that accentuatesisotopic fractionation. Highd11B also reflects the lack of asignificant thickness of isotopically negative, continentally-derived sediments along much of the Tonga Trench.

7.5. Influence of Louisville Ridge

Subduction and southward migration of the Louisville Ridgealong the Tonga forearc (Herzer and Exon, 1985) might havebeen expected to increase sediment subduction due to thepresence of its sediment apron and as a result of the tectonicerosion and subduction of volcaniclastic sequences within theforearc. Louisville Ridge is a Cretaceous-Late Cenozoic hot-spot related feature (Lonsdale, 1988) with more than 60 vol-

3358 Peter D. Clift et al.

canoes arranged at intervals,100 km apart along a 75-km-wide band stretching;4300 km towards the SE from Tonga.The Tonga Trench is in collision with Osborne Seamount justsouth of Ata (Lonsdale, 1986). On the scale of the studypresented here Louisville Ridge may be considered as a pointsource of material flux to the trench. Packham (1985) estimatedthat the Louisville Ridge began to collide with the northern-most part of Tonga forearc shortly after 4 Ma, and sweptrapidly down the margin to its present location. Nonetheless,there is no lowd11B measured at Tafahi, as might be expectedfor the involvement of larger volume of sediment with negativeboron isotopic composition in petrogenesis.

The Pb and Sr isotopic character of Tafahi and Niuatoputapuvolcanic rocks has been used to argue that Louisville RidgeAOC or volcaniclastic sediment was influential over petrogen-esis in this area (Ewart et al., 1998; Regelous et al., 1997;Turner and Hawkesworth, 1997). Figures 8B and C show thealong-strike variation in230Th/232Th and 87Sr/86Sr, as mea-sured by Regelous et al. (1997) and Ewart et al. (1998). Theincreasing87Sr/86Sr and decreasing230Th/232Th correlate withfalling d11B in the Kermadec Arc and support a relationshipbetween the proportion of sediment in the slab flux andd11Bthere. Comparison with Indian and Pacific mantle values (87Sr/86Sr,0.7034) suggests that the Sr system in the Kermadec Arcis strongly contaminated by sediment, whereas decreasing87Sr/86Sr moving northward along the Tonga Arc can be related tocontamination by the lavas and volcaniclastic sediment of theplume-derived Louisville Ridge.

However, at the north end there is no correlation in theTonga Arc between the Sr and Th, andd11B. Our resultsuggests that the boron isotopic characteristics of the Louisvillevolcaniclastic apron differ from the continental sediments inthe Kermadec Trench and are similar to the AOC of theLouisville Ridge itself. Given that the ultimate source of theboron in the Pacific AOC and the Louisville Ridge AOC is thesame, i.e., seawater, it is not surprising that dewatering ofLouisville Ridge does not affect the boron isotopic character ofthe volcanic rocks at Tafahi and Niuatoputapu. In contrast,Pacific AOC and Louisville AOC have totally different Sr, Thand Pb isotope compositions (Regelous et al., 1997), so that theSr, Th and Pb isotope systematics of these islands are disruptedby ridge subduction.

7.6. Temporal Variations in Fluid Source

Typically highd11B values at ODP Site 840 suggest that slabflux has been dominated by water expelled from the subductingAOC at least in the region of Ata since 7 Ma. Brief excursionsto low d11B values are noted, with some sediments showinghuge internal variations ind11B values. These sediments incor-porate material from several eruption events, possibly incorpo-rating material transported far along the arc or reworked fromolder deposits, since the degree of variability would be hard toaccount for in a single eruption. Short-term variations ind11Bare likely due to changes in the proportion of sediment andAOC-derived fluids. This is because the lag time required toheat or cool a subduction zone following changes in the age ofthe subducting slab and its rate of convergence is too long toaffect the boron isotope character of the arc over short timescales (Peacock, 1996). Given the composition of the modern

Tonga Arc, the large distance of ODP Site 840 from NewZealand (and thus probable lack of input from sediments de-rived from there) and the apparent inability of volcaniclasticsediment to strongly alter thed11B values of the arc, thistemporal history is not surprising. The relative lack of short-term variations ind11B contrasts with the conclusions of Ish-ikawa and Nakamura (1999) in the Marianas Arc. We suggestthat the relatively featureless character of the subducting platein Tonga, compared to the seamount strewn Pacific east of theMarianas, results in a more constant supply of water dominatedby the AOC.

8. INFLUENCE OF FLUID FLUX ON PETROGENESIS

We now examine the coherence of variations in the slabfluid-flux with other elemental groups to determine the effectthat fluid-flux has on the arc chemistry. Slab fluid-flux to thebackarc is generally significantly less than that to the main arcvolcanic front (Pearce et al., 1995). Consequently the volcanicfront is the best place to examine the influence of fluid flux, ascharted by boron, over other element groups. Clift and Dixon(1994) noted an up-section trend to decreasing alkali elementconcentrations from a maximum at 7 Ma, to a minimum at 3 to4 Ma. K and Na show good negative correlation with B/Be, i.e.,the volume of slab fluid-flux (Fig. 9. Contrary to their behaviorin many arcs the alkali elements behave dominantly as incom-patible elements and do not show enrichment by fluid-flux fromthe slab. Instead the falling maximum values reflect eitherincreasing degrees of depletion in the mantle wedge or increas-ing degrees of melting of a relatively constant mantle source.Increased fluid flux can account for the trend by increasingpartial melting through lowering of the solidus at any givendepth in the wedge. The slab-derived fluids cannot have con-tained large amounts of alkali elements, despite their mobilityin water. The K2O content of deep-sea clays is;3% (e.g.,Underwood et al., 1993), compared to 0.1 to 0.2% in mid oceanridge basalts (Melson et al., 1976). The low concentration ofthe alkali elements in the arc lavas is thus consistent with themajority of the slab fluid-flux being derived from the AOC.

To minimize the effects of fractional crystallization on theimage of mantle melting conditions provided by the arc lavasand tephras we chose to examine the relative enrichment ofHFSEs and REEs in only the more basaltic tephra shards (SiO2

,60%) from the ODP wells. Figure 10 shows the relationshipbetween Nb/Zr and La/Sm, and B/Be. Nb/Zr and La/Sm may beconsidered proxies for the relative enrichment in the HFSEsand REEs respectively. Both groups show highest depletionwhen B/Be is high (i.e., at 3–4 Ma), although there is signifi-cant scatter. Such a relationship is in accord with the alkalielement data (Fig. 9) in suggesting higher degrees of partialmelting when slab fluid-flux is high. Moreover, the positivecorrelation of B/Be and both Nb/Zr and La/Sm is consistentwith the measured B/Be being magmatic and not substantiallyaffected by diagenesis.

We propose an important modification to models that linkREEs and slab flux, notably those in which LREE enrichmentis dominated by sediment dewatering or melting (e.g., Plank etal., 1993; Hole et al., 1984). Instead our new data demonstrateLREE depletion driven by dewatering mostly of the AOC, asshown by high B/Be andd11B values in the volcaniclastic

3359Boron isotopic tracing of slab-flux in Tonga

Fig. 8. (A) Diagram showing the variations in silica versus the percentage of volatiles at ODP Site 840, estimated by the100%-total measured major element composition. A positive correlation of SiO2 and volatiles is a natural progression of thefractionation process. Those grains with anomalously high volatile content, due to alteration can then be identified. (B) Thetemporal variation in volatiles shows that the vast majority of the altered grains were deposited before 5 Ma. (C) Nocorrelation is noted between boron isotopic ratios and degree of volatile content. (D) Some of the least altered grains showvery highd11B values suggesting that these are original and magmatic values.

3360 Peter D. Clift et al.

shards. The very positived11B and high B/Be values measuredargue against slab melting in a hot subduction, because thatwould prevent the B isotope fractionation required here (Pea-cock and Hervig, 1999). In areas where slab melting is knownto be important (e.g., western Aleutians, Yogodzinski andKelemen, 1998; Cascades, Hughes, 1990), B/Be is very low(5.9–39 in Aleutians, 3.6–4.7 in Cascades; Morris et al., 1990).

It is noteworthy that LREE enrichment does not correlatewith the peak time of sediment subduction (;2–3 Ma) noted byClift and Vroon (1996). Even if the temporal correlation isaccepted, peak sediment flux might be expected to cause LREEenrichment, not the depletion seen at that time, if the sediment

chemical signal had been transferred to the arc lavas by melt.The volcaniclastic record suggests that variations in the volumeof fluid fluxed from the subducting slab, and tracked by B/Be,can explain the chemical variability by controlling the degreeof partial melting, since addition of water to the source mantlewill depress the solidus. In practice this is the fluid-flux meltinghypothesis of Luhr (1992), Ryan and Langmuir (1993) andStolper and Newman (1994). Explanations of the chemicalvariability based on the rifting tectonics of the Lau Basincontrolling the height of the melting column (e.g., Clift andDixon, 1994) are consistent with the LREE and HFSE data butdo not account for the B/Be variability. Temporal variations inslab fluid-flux alone are capable of reconciling all three chem-ical groups.

9. CONCLUSIONS

We demonstrate here that since 7 Ma there have been sig-nificant changes in the relative volume of slab fluid-flux intothe Tonga Arc in the vicinity of Ata, and that these variationshave been much higher than any variability now seen along thestrike of the Tonga-Kermadec Arc. Peak B/Be at 3 to 4 Ma may

Fig. 9. Variations in degree of incompatible element depletion in (A)the HFSEs (Nb/Zr) and (B) REEs (La/Sm) with B/Be for single-grainbasaltic andesite shards.-

Fig. 10. Diagram showing the temporal variations in total K2Ocontents for glass shards of all compositional ranges at ODP Site 834and 840. Note the minimum values at 3 to 4 Ma. Data from Clift andDixon (1994).

3361Boron isotopic tracing of slab-flux in Tonga

represent either the effect of accelerated subduction due toregional plate readjustments at;5 Ma (Cande et al., 1995) orthe influence of Lau Basin rifting and the start of slab roll-backtowards the east. A regional explanation of increased westwardplate velocity would account for the peak in B/Be also seen inthe Marianas at 3 to 4 Ma. The high peak B/Be values are notcompatible with simple contamination of the mantle source bysediment, and require much of the boron to be derived byaqueous fluid from the AOC, an explanation compatible withthe high positived11B compositions. However, extreme isoto-pic fractionation within a cold, fast subduction zone is alsorequired to explain the very highd11B values observed. Peaksin the volume of slab fluid-flux measured by B/Be do not reflectperiods of preferential sediment subduction, but may indicatechanges in the rate of plate convergence. The collision ofLouisville Ridge with the Tonga Trench does not seem to affectthe boron systematics.

Latitudinal variation in the boron isotopic composition alongthe Tonga-Kermadec Arc can be linked to greater degrees ofsediment involvement in petrogenesis close to New Zealand,but not to passage of the Louisville Ridge in northern Tonga.This pattern reflects the distinctive negatived11B compositionof the continental-derived material, but the similarity of thepositive boron isotopic character of the basement and volcani-clastic cover of the Louisville Ridge to with normal oceaniccrust.

Periods of increased slab flux (high B/Be) correlate withgreater depletion in HFSEs and REEs, and suggest that flux ofaqueous fluid from the AOC is the primary control on thedegree of partial melting under the arc.

Acknowledgments—We wish to thank Tony Ewart and David Tappinfor their donation of modern volcanic specimens from the Tonga-Kermadec Arc. PC thanks Bill Bryan, Jim Hawkins and Tony Ewart forfirst interesting him in the Tonga Arc. Support for EFR was providedby French “Programme Lavoisier” Post-Doctoral Fellowship and J.Steward Johnson Scholar from WHOI. We thank Achim Kopf andMarc Chaussidon for their helpful reviews in improving this paper.This is WHOI contribution 10502.

Associate editor:D. B. Dingwell

REFERENCES

Austin J. A. Jr., Taylor F. W. and Cagle C. D. (1989) Seismic stratig-raphy of the central Tonga Ridge.Mar. Petrol. Geol.6, 71–92.

Bebout G. E., Ryan J. G., Leeman W. P., and Bebout A. E. (1999)Fractionation of trace elements by subduction-zone metamorphism;effect of convergent-margin thermal evolution.Earth Planet. Sci.Lett. 171,63–81.

Brown L., Klein J., Middleton R., Sacks I. S., and Tera F. (1982) 10Bein island arc volcanoes and implications for subduction.Nature299,718–720.

Burnham C. W. and Jahns R. H. (1962) A method of determining thesolubility of water in silicate melts.Am. J. Sci.260,721–745.

Burns R. E., Andrews J. E., et al. (1973) Site 204.Init. Rpts. Deep SeaDrill. Proj. 21, 33–56.

Cande S. C., Raymond C. A., Stock J., and Haxby W. F. (1995)Geophysics of the Pitman fracture zone and Pacific-Antarctic platemotions during the Cenozoic.Science270,947–953.

Chaussidon M., Robert F., Mangin D., Hanon P., and Rose E. (1997)Analytical procedures for the measurement of boron isotope com-positions by ion microprobe in meteorites and mantle rocks.J.Geostand. Geoanal.21, 7–17.

Chaussidon M. and Marty B. (1995) Primative boron isotope compo-sition of the mantle.Science269,383–386.

Chaussidon M. and Libourel G. (1993) Boron partitioning in the uppermantle; an experimental and ion probe study.Geochim. Cosmochim.Acta 57, 5053–5062.

Clift P. D., and MacLeod C. J. (1999) Slow rates of tectonic erosionestimated from the subsidence and tilting of the Tonga ForearcBasin.Geol.27, 411–414.

Clift P. D., and Blusztajn J. (1999) The Influence of Slab Flux onPetrogenesis in the Aegean and Aeolian Volcanic Arcs tracedthrough Microprobe Analysis of Marine Tephras.J. Vol. Geotherm.Res.92, 321–347.

Clift P. D. and Lee J. (1998) Temporal evolution of the Mariana Arcduring rifting of the Mariana Trough traced through the volcaniclas-tic record.The Island Arc.7, 496–512.

Clift P. D., MacLeod C. J., Tappin D. R., Wright D., and Bloomer S. H.(1998) Tectonic controls on sedimentation in the Tonga Trench andforearc, SW Pacific.Bull. Geol. Soc. Am.110,483–496.

Clift P. D. and Vroon P. Z. (1996) Isotopic evolution of the Tonga Arcsystem during Lau Basin rifting; Evidence from the volcaniclasticrecord.J. Petrol.37, 1153–1173.

Clift P. D. and ODP Leg 135 Scientific Party. (1995) Volcanism andSedimentation in a rifting island arc terrain; an example from Tonga,SW Pacific. InVolcanism Associated with Extension at ConsumingPlate Margins(ed. J. L. Smellie),Geol. Soc., London, spec. publ.88,29–52.

Clift P. D. and Dixon J. E. (1994) Variations in arc volcanism andsedimentation related to rifting of the Lau Basin, SW Pacific.Proc.Ocean Drill. Prog., Sci. Res.135,23–50.

Clift P. D. (1994) Controls on the sedimentary and subsidence historyof an active plate margin: an example from the Tonga Arc, SWPacific.Proc. Ocean Drill. Prog., Sci. Res.135,173–189.

DeMets C., Gordon R. G., Argus D. F. and Stein S. (1990) Currentplate motions.Geophys. J. Int.101,425–478.

Edwards C. M. H., Morris J. D. and Thirlwall M. F. (1993) Separatingmantle from slab signatures in arc lavas using B/Be and radiogenicisotope systematics.Nature362,530–533.

Elliott T., Plank T., Zindler A., White W., and Bourdon B. (1997)Element transport from slab to volcanic front at the Mariana Arc.J.Geophys. Res.102,14,991–15,019.

Ewart A., Collerson K. D., Regelous M., Wendt J. I., and Niu Y. (1998)Geochemical evolution within the Tonga-Kermadec-Lau Arc-back-arc systems: the role of varying mantle wedge composition in spaceand time.J. Petrol.39, 331–368.

Gaillardet J. and Allegre C. J. (1995) Boron isotopic compositions ofcorals; seawater or diagenesis record?Earth Planet. Sci. Lett.136,665–676.

Gill J. B. (1976) Composition and age of Lau Basin and Ridge volcanicrocks: Implications for evolution of an inter-arc basin and remnantarc.Bull Geol. Soc. Am.87, 1384–1395.

Gill J. B., Morris J. D. and Johnson R. W. (1993) Timescale forproducing the geochemical signature of island arc magmas; U-Th-Poand Be-B systematics in Recent Papua New Guinea lavas.Geochim.Cosmochim. Acta57, 4269–4283.

Hawkins J. W., Castillo P. R., and Lonsdale P. (1999) Petrology of theoceanic crust entering the Kermadec Trench between 28 and 32°S.EOS,trans.80, 1158.

Hawkins J. W. (1974) Geology of the Lau Basin, a marginal sea behindthe Tonga Arc. InGeology of Continental Margins(eds. C. Burkeand C. Drake), pp. 505–520, Springer Verlag, Berlin.

Hemming N. G. and Hanson G. N. (1992) Boron isotopic compositionand concentration in modern marine carbonates.Geochim. Cosmo-chim. Acta56, 537–543.

Herzer R. H. and Exon N. F. (1985) Structure and basin analysis of thesouthern Tonga forearc. InGeology and Offshore Resources of thePacific Island Arcs-Tonga Region(eds. D. W. Scholl and T. L.Vallier), Circum-Pac. Counc. Energy Miner. Resour.,Earth Sci. Ser.2, 55–73.

Hiscott R. N. and Gill J. B. (1992) Major and trace element geochem-istry of Oligocene to Quaternary volcaniclastic sands and sandstonesfrom the Izu-Bonin Arc.Proc. Ocean Drill. Prog., Sci. Res.126,467–486.

Hochstaedter A. G., Ryan J. G., Luhr J. F., and Hasenaka T. (1996) OnB/Be ratios in the Mexican volcanic belt.Geochim. Cosmochim.Acta 60, 613–628.

3362 Peter D. Clift et al.

Hole M. J., Saunders A. D., Marriner G. F., and Tarney J. (1984)Subduction of pelagic sediment: implications for the origin of Ce-anomalous basalts from the Mariana Islands.J. Geol. Soc. London.141,453–472.

Hughes S. S. (1990) Mafic magmatism and associated tectonism of thecentral high Cascade Range, Oregon.J. Geophys. Res.95, 19,623–19,638.

Ishikawa T. and Tera F. (1999) Two isotopically distinct fluid compo-nents involved in the Mariana Arc; evidence from Nb/B ratios and B,Sr, Nd, and Pb isotope systematics.Geol.27, 83–86.

Ishikawa T. and Nakamura E. (1994) Origin of the slab component inarc lavas from across-arc variation of B and Pb isotopes.Nature370,205–208.

Ishikawa T. and Nakamura E. (1993) Boron isotope systematics ofmarine sediments.Earth Planet. Sci. Lett.117,567–580.

Johnson M. C. and Plank T. (1999) Dehydration and melting experi-ments constrain the fate of subducted sediments.Geochem. Geophys.Geosyst.1.

Karig D. E. (1970) Ridges and basins of the Tonga-Kermadee islandarc system.J. Geophys. Res.75, 239–254.

Kay R. W., Sun S. S. and Lee-Hu C. N. (1978) Pb and Sr isotopes involcanic rocks from the Aleutian Islands and Pribilof Islands,Alaska.Geochim. Cosmochim. Acta42, 263–274.

Leeman W. P., Carr M. J., and Morris J. D. (1994) Boron geochemistryof the Central American volcanic arc; constraints on the genesis ofsubduction-related magmas.Geochim. Cosmochim. Acta58, 149–168.

Lin P. N., Stern R. J., Morris J., and Bloomer S. H. (1990) Nd- and Sr-isotopic compositions of lavas from the northern Mariana and south-ern Volcano arcs: implications for the origin of island arc melts.Contrib. Mineral. Petrol.105,381–392.

Lonsdale P. F. (1988) Geography and history of the Louisville hotspotchain in the Southwest Pacific.J. Geophys. Res.93, 3078–3104.

Lonsdale P. F. (1986) A multibeam reconnaissance of the TongaTrench axis and its intersection with the Louisville guyot chain.Mar.Geophys. Res.8, 295–327.

Luhr J. F. (1992) Slab-derived fluids and partial melting in subductionzones; insights from two contrasting Mexican volcanoes (Colimaand Ceboruco).J. Vol. Geotherm. Res.54, 1–18.

McMullen C. C., Cragg C. B. and Thode H. G. (1961) Absolute ratioof 11B/10B in Searles Lake borax.Geochim. Cosmochim. Acta23,147–150.

Melson W. G., Vallier T. L., Wright T. L., Byerly G. and Nelen J.(1976) Chemical diversity of abyssal volcanic glass erupted alongPacific, Atlantic, and Indian Ocean sea-floor spreading centersAm.Geophys U., Geophys. Monog.19, 351–367.

Miller D. M., Goldstein S. L. and Langmuir C. H. (1994) Cerium/leadand lead isotope ratios in arc magmas and the enrichment of lead inthe continents.Nature368,514–520.

Moriguti T. and Nakamura E. (1998) Across-arc variation of Li iso-topes in lavas and implications for crust/mantle recycling at subduc-tion zones.Earth Planet. Sci. Lett.163,167–174.

Moran A. E., Sisson V. B., and Leeman W. P. (1992) Boron depletionduring progressive metamorphism; implications for subduction pro-cesses.Earth Planet. Sci. Lett.111,331–349.

Morris J. D., Ryan J. and Leeman W. P. (1993) Be isotope and B-Beinvestigations of the historic eruptions of Mt. Vesuvius.J. Volcan.Geotherm. Res.58, 345–358.

Morris J. D., Leeman W. P., and Tera F. (1990) The subductedcomponent in island arc lavas: constraints from Be isotopes andBe-B systematics.Nature344,31–36.

Oi T., Nomura M., Musashi Ma., Ossaka T., Okamoto M. and Kaki-hana H. (1989) Boron isotopic compositions of some boron minerals.Geochim. Cosmochim. Acta53, 3189–3195.

Packham G. H. (1985) Vertical tectonics on the Tonga Ridge from theTongatapu oil exploration wells. InGeology and Offshore Resourcesof the Pacific Island Arcs-Tonga Region(eds. D. W. Scholl and T. L.Vallier), Circum-Pac. Counc. Energy Miner. Resour.,Earth Sci. Ser.2, 291–300.

Palmer M. R. (1991) Boron-isotope systematics of Halmahera Arc(Indonesia) lavas; evidence for involvement of the subducted slab.Geol.19, 215–217.

Palmer M. R., Spivack A. J. and Edmond J. M. (1987) Temperature and

pH controls over isotopic fractionation during adsorption of boron onmarine clay.Geochim. Cosmochim. Acta51, 2319–2323.

Parson L. M., Hawkins J. W., Allan J. F, et al. (1992) Proceedings ofthe Ocean Drilling Program, Lau Basin covering Leg 135 of thecruises of the drilling vessel JOIDES Resolution, Suva Harbor, Fijiho Honolulu, Hawaii, site 834-841, Dec. 1990–28 Feb. 1991Proc.Ocean Drill. Prog. Init. Res.135,0–643.

Peacock S. M., and Hervig R. L. (1999) Boron isotopic composition ofsubduction-zone metamorphic rocks.Chem. Geol.160,281–290.

Peacock S. M. and Hervig R. L. (1997) Boron isotopes and the transferof subducting slab components into arc magmas.Eos,trans.78,840.

Peacock S. M. (1996) Thermal and petrologic structure of subductionzones. InSubduction top to bottom(eds. G. E. Bebout, D. W. Scholl,S. H. Kirby and J. P. Platt),Am. Geophys. U., Geophys. Monog.96,119–133.

Pearce J. A. (1983) Role of the sub-continental lithosphere in magmagenesis at active continental margins. InContinental basalts andmantle xenoliths(eds. C. J. Hawkesworth and M. J. Norry) pp.230–249. Shiva, Chichester, geol. ser.

Pearce J. A., Ernewein M., Bloomer S. H., Parson L. M., Murton B. J.,and Johnson L. E. (1995) Geochemistry of Lau Basin volcanic rocks:influence of ridge segmentation and arc proximity. InVolcanismAssociated with Extension at Consuming Plate Margins(ed. J. L.Smellie)Geol. Soc., London, Spec. Publ.88, 53–75.

Pelletier B., Calmant S., and Pillet R. (1998) Current tectonics of theTonga-New Hebrides region.Earth Planet. Sci. Lett.164,263–276.

Plank T. and Langmuir C. H. (1998) The chemical composition ofsubducting sediment: implications for the crust and mantle.Chem.Geol.145,325–394.

Plank T. and Langmuir C. H. (1993) Tracing trace elements fromsediment input to volcanic output at subduction zones.Nature362,739–743.

Plank T. and Langmuir C. H. (1988) An evaluation of the globalvariations in the major element chemistry of arc basalts.EarthPlanet. Sci. Lett.90, 349–370.

Reagan M. K., Morris J. D., Herrstrom E. A. and Murrell M. T. (1994)Uranium series and beryllium isotope evidence for an extendedhistory of subduction modification of the mantle below Nicaragua.Geochim. Cosmochim. Acta.58, 4199–4212.

Regelous M., Collerson K. D., Ewart A., and Wendt J. I. (1997) Traceelement transport rates in subduction zones:evidence from Th, Sr andPb isotope data for Tonga-Kermadec arc lavas.Earth Planet. Sci.Lett. 150,291–302.

Rose E. F. (1999) Geochimie isotopique du bore dans les cylcessupergenes. Ph.D. Thesis, Institut Nationale Polytechnique de Lor-raine, Nancy, France.

Ryan J. G. and Langmuir C. H. (1993) The systematics of boronabundances in young volcanic rocks.Geochim. Cosmochim. Acta57,1489–1498.

Ryan J. G. (1989) The systematics of lithium, beryllium and boron inyoung volcanic rocks, Ph.D. thesis, Columbia University, New York,0–326.

Seyfried W. E. Jr, Janecky D. R. and Mottl M. J. (1984) Alteration ofthe oceanic crust; implications for geochemical cycles of lithium andboron.Geochim. Cosmochim. Acta48, 557–569.

Shipboard Scientific Party (1992) Site 834.Proc. Ocean Drill. Prog.Init. Res.135,85–180.

Smith H. J., Leeman W. P., Davidson J. and Spivack A. J. (1997) TheB isotopic composition of arc lavas from Martinique, Lesser Anti-lles. Earth Planet. Sci. Lett.146,303–314.

Smith H. J., Spivack A. J., Staudigel H., and Hart S. R. (1995) Theboron isotopic composition of altered oceanic crust.Chem. Geol.126,119–135.

Sobolev A. V. and Chaussidon M. (1996) H2O concentrations inprimary melts from supra-subduction zones and mid-ocean ridges;implications for H2O storage and recycling in the mantle.EarthPlanet. Sci. Lett.137,45–55.

Spivack A. J. and Edmond J. M. (1987) Boron isotope exchangebetween seawater and the oceanic crust.Geochim. Cosmochim. Acta51, 1033–1043.

Spivack A. J., Palmer M. R. and Edmond J. M. (1987) The sedimentarycycle of the boron isotopes.Geochim. Cosmochim. Acta51, 1939–1949.

3363Boron isotopic tracing of slab-flux in Tonga

Stolper E. and Newman S. (1994) The role of water in the petrogenesisof Mariana Trough magmas.Earth Planet. Sci. Lett.121,293–325.

Sun S.-S. and McDonough W. F. (1989) Chemical and isotopic sys-tematics of oceanic basalts: implications for mantle composition andprocesses. InMagmatism in the Ocean Basins(eds. A. D. Saundersand M. J. Norry),Geol. Soc. Lond., Spec. Publ.42, 313–346.

Tagudin J. E. and Scholl D. W. (1994) The westward migration of theTofua volcanic arc towards the Lau Basin. InGeology and Resourcesof Island Arcs—Tonga-Lau-Fiji Region(eds. R. H. Herzer, P. F.Ballance and A. J. Stevenson)SOPAC Tech. Bull.8, 121–130.

Tappin D. R., Bruns T., Geist E. L., and Lavoie D. (1994) Correlationof regional seismic stratigraphy with the sedimentary sequence atSite 840.Proc. Ocean Drill. Prog., Sci. Res.135,331–366.

Tatsumi Y. (1989) Migration of fluid phases and genesis of basaltmagmas in subduction zone.J. Geophys. Res.94, 4697–4707.

Tatsumi Y., and Isoyama H. (1988) Transportation of beryllium withH2O at high pressures; implication for magma genesis in subductionzones.Geophys. Res. Lett.15, 180–183.

Tatsumi Y., Sakuyama M., Fukuyama H., and Kushiro I. (1983) Gen-eration of arc basalt magmas and thermal structure of the mantlewedge in subduction zones.J. Geophys. Res.88, 5815–5825.

Tera F., Brown L., Morris J., Sacks S., Klein J. and Middleton R.(1986) Sediment incorporation in island arc magmas: inferencesfrom 10Be.Geochim. Cosmochim. Acta50, 535–550.

Thompson G. and Melson W. G. (1970) Boron contents of serpentinitesand metabasalts in the oceanic crust; implications for the boron cyclein the oceans.Earth Planet. Sci. Lett.8, 61–65.

Turner S., Hawkesworth C., Rogers N., Bartlett J., Worthington T.,Hergt J., Pearce J., and Smith I. (1997)238U-230Th disequilibria,magma petrogenesis and flux rates beneath the depleted Tonga-Kermadec island arc.Geochim. Cosmochim. Acta61, 4855–4884.

Turner S. and Hawkesworth C. (1997) Constraints on flux rates andmantle dynamics beneath island arcs from Tonga-Kermadec lavageochemistry.Nature389,568–573.

Underwood M. B., Pickering K., Gieskes J. M., Kastner M. and Orr R(1993) Sediment geochemistry, clay mineralogy, and diagenesis; asynthesis of data from Leg 131, Nankai Trough.Proc. Ocean Drill.Prog., Sci. Res.131,343–363.

Yogodzinski G. M. and Kelemen P. B. (1998) Slab melting in theAleutians; implications of an ion probe study of clinopyroxene inprimitive adakite and basalt.Earth Planet. Sci. Lett.158,53–65.

You C. F., Spivack A. J., Gieskes J. M., Rosenbauer R., and BischoffJ. L. (1995) Experimental study of boron geochemistry; implicationsfor fluid processes in subduction zones.Geochim. Cosmochim. Acta59, 2435–2442.

You C. F., Spivack A. J., Smith J. H., and Gieskes J. M. (1993)Mobilization of boron in convergent margins; implications for theboron geochemical cycle.Geol.21, 207–210.

3364 Peter D. Clift et al.