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Geochimica et Cosmochimica Acta, Vol. 65, No. 19, pp. 3347–3364, 2001 Copyright © 2001 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/01 $20.00 ⫹ .00 Pergamon PII S0016-7037(01)00670-6 Tracing the evolving flux from the subducting plate in the Tonga-Kermadec arc system using 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 REGELOUS2 1 Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA 2 Max-Planck-Institut fur Chemie, Abteilung Geochemie, Postfach 3060, 55020 Mainz, Germany (Received July 24, 2000; accepted in revised form May 4, 2001) Abstract—The influence of fluid flux on petrogenesis in the Tonga-Kermadec Arc was investigated using ion microprobe measurements of B/Be and boron isotope ratios (11B/10B) to document the source and relative volumes of the fluids released from the subducting oceanic plate. We analyzed young lavas from eight different islands along the Tonga-Kermadec Arc, as well as glass shards in volcanic sediments from Ocean Drilling Program (ODP) Site 840, which record the variations in the chemistry of Tonga magmatism since 7 Ma. B/Be is variable (5.8 –122), in young Tonga-Kermadec Arc lavas. In contrast, glass shards from ⬃3 to 4 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 very high ␦11B values (⫺11.6 to ⫹37.5) for the same shards and lavas these data indicate that most of the B is derived from fluid escaped from the subducting altered Pacific oceanic crust, rather than from sediment. High ␦11B values also reflect large degrees of isotopic fractionation in this cold fast subduction zone. Lower ␦11B values noted in the Kermadec Arc (17 to ⫺4.4) are related to the influence of sediment eroded from New Zealand and slower convergence. High fluid flux (B/Be) is synchronous in Tonga and the Marianas at 3 to 4 Ma 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 field strength element depletion. This suggests maximum degrees of partial melting at this time. Although thinning of 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 trace element systematics in the Tonga-Kermadec Arc with no indication of slab melting. Copyright © 2001 Elsevier Science Ltd the altered oceanic crust (AOC) and the sedimentary cover (Stolper and Newman, 1994), and those models in which the sedimentary cover dominates (Hole et al., 1984; Elliott et al., 1997; Plank and Langmuir, 1998). Different methods have been used to demonstrate slab involvement in arc magmatism. The high concentrations of large ion lithophile elements (LILE) relative to high field strength (HFSE) and rare earth elements (REE) in arc lavas, compared to midocean ridge basalts, show that the LILE budget of arc lavas is typically derived from subducted materials (e.g., Pearce, 1983). For example, oceanic basalts have uniform Ce/Pb ratios of about 25 (Sun and McDonough, 1989), whereas most arc lavas have much lower Ce/Pb, due to addition of Pb from subducted sediment and altered oceanic crust (e.g., Miller et al., 1994). Similarly, much of the Ba, U, Rb, Sr, as well as B, in arc lavas is also derived from the subducted slab. Isotopic data also demonstrate slab involvement in arc magmatism. The cosmogenic isotope 10Be has been found in several arcs (e.g., Brown et al., 1982; Tera et al., 1986; Morris et al., 1990), implying rapid sediment subduction and melting, because this element is radioactive with a half-life of only 1.5 m.y. Similarly Sr and Pb isotopes, especially the 207Pb/ 204 Pb ratio, has been used as strong evidence for sediment involvement in several arcs (e.g., Aleutians and Alaska; Kay et al., 1978). However, the influence of the altered oceanic crust (AOC), as well as the sediments, has also been highlighted using Pb isotopes (Miller et al., 1994) and the boron system (Spivack and Edmond, 1987; Seyfried et al., 1984). 1. INTRODUCTION Understanding the processes that control arc petrogenesis is one of the outstanding problems in modern petrology, but a difficult task due to the large number of potential variables compared to the relatively straight forward decompression melting of a mid ocean ridge system. In most subduction systems the arc volcanic front is located over the dewatering slab at around 100 km depth (Tatsumi et al., 1983). The fluid and amphibole, generated by hydration of the mantle wedge above a subducting oceanic plate, is carried down and under the arc due to convection in the wedge. As it does so the material is warmed and melting begins to occur above about 1000°C (Tatsumi et al., 1983). Further melting occurs due to decompression during asthenospheric upwelling in a melt column under the arc volcanoes. Petrogenetic models for subduction systems have focused on two key variables in controlling petrogenesis. In one set of models, typified by Plank and Langmuir (1988), lithospheric processes dominate in constraining the extent of partial melting by limiting the degree of upwelling of the mantle wedge above the subducting plate. In the other set of models the geochemical character of an arc’s output is mostly controlled by the nature of the fluid flux from the subducting plate. Within this latter group of models two approaches have been taken, those in which the fluid is variable and comprises of components from *Author to whom correspondence should be addressed (pclift@ whoi.edu). 3347 3348 Peter D. Clift et al. If fluid flux is the principal control on petrogenesis then variations in this flux should have a major effect on arc geochemistry. Hole et al. (1984), and more recently Plank and Langmuir (1993) and Johnson and Plank (1999), suggested that the REE characteristics of many arc lavas are derived from the sediments being subducted in the adjacent trench. In the Marianas a relative negative anomaly in Ce seen in the lavas from Agrigan Island was compared with a negative Ce anomaly in the sediments on the Pacific Plate adjacent to this section of the arc. Indeed, Hole et al. (1984) suggested that the light REE (LREE) enrichment known from the Northern Marianas may reflect subduction of a large thickness of hotspot-derived clastics, rather than a long lived chemical anomaly in the mantle, as proposed by Lin et al. (1990). If hypotheses such as these are correct then understanding the nature and controls on flux from the subducting oceanic slab is crucial to understanding the extent to which subducted sediment is recycled back into the mantle, and the process of new crust formation at convergent plate margins. In this paper we examine a single key variable in this system, the fluid flux from the subducting slab within the classic intraoceanic Tonga subduction system using the boron system. We exploit the temporal dimension of arc volcanic activity available through the volcaniclastic record of the forearc basin to chart the evolution in slab flux to the arc volcanic front in terms of its magnitude and origin, and examine how this has changed since 7 Ma. 2. SETTING OF TONGA ARC The volcanically active Tofua Arc and Lau Basin represent a classic arc and backarc basin system (Karig, 1970; Hawkins, 1974; Gill, 1976), which has developed as a result of subduction of the Pacific lithosphere under the eastern edge of the Indian-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 additional spreading in the Lau Basin (Pelletier et al., 1998). Subduction commenced in this region at approximately 45 Ma (Herzer and Exon, 1985) since which time the plate margin appears to have been in a constant state of subduction erosion (Clift and MacLeod, 1999). This means that sediment on the Pacific Plate has not accreted during some periods and then been eroded later to produce pulses in sediment flux to the arc, but has instead been always fully subducted, along with varying amounts of the forearc basement and sedimentary cover. The present arc system dates from the Late Miocene, when the Lau Basin was generated by rifting of the Tonga Arc (Parson et al., 1992). Rifting had started by at least 7.0 Ma, although the generation of new backarc crust may postdate this by as much 1.4 m.y. A minimum age of 5.6 Ma is given for the oldest igneous crust in the new basin at ODP Site 834 (Fig. 1; Shipboard Scientific Party, 1992). The original Miocene arc is now preserved on the western side of the Lau Basin as a remnant arc called the Lau Ridge. Activity on the new “Tofua Arc” dates from about 3.0 Ma (middle Pliocene; Tappin et al., 1994; Clift et al., 1995). The Tonga Platform has thus been in a forearc position to both the Miocene Tonga and the new Tofua Arcs, and so in a position to record the volcanic products of each. Volcanism in the Lau Basin is not thought to have Fig. 1. Bathymetric map of the Tonga-Kermadec Arc showing the location of the principle bathymetric features, scientific drill sites and active volcanic centers described in this study. ELSC ⫽ Eastern Lau Spreading Center, CLSC ⫽ Central Lau Spreading Center. contributed to the sedimentation on the Tonga Platform, due to the great difference in water depths and the submarine character of the volcanism in that area (Clift et al., 1995). 3. PREVIOUS WORK The Tonga-Kermadec Arc system has been the target of many geochemical studies. Pearce et al. (1995) highlighted the subduction input by showing that basalts from the Eastern Lau Spreading Center have much higher enrichment in LILEs relative to the Central Lau Spreading Center (Fig. 1). Trace element and Th, Sr and Pb isotopic data from the modern arc volcanic centers have been used to argue that in northern Tonga much of the slab flux is from the altered oceanic crust, contrasting with the more sediment-dominated signature seen in the Kermadec Arc to the south, where a greater thickness of Boron isotopic tracing of slab-flux in Tonga sediment is being subducted (Regelous et al., 1997; Ewart et al., 1998). Clift and Vroon (1996) used Pb, Nd and Sr isotopic data from the volcaniclastic record to support the idea of a temporary peak in the subducted sediment contribution to the arc following the start of backarc rifting. There was a delay of ⬃3 m.y. between the onset of rifting and the peak of the recognized sediment pulse, a figure that agreed well with the conclusions of Regelous et al. (1997), Turner and Hawkesworth (1997) and Turner et al. (1997) that 2 to 4 m.y. is required for Pb in the material subducted at the trench to be erupted at the adjacent volcanic front. Clift and Dixon (1994) analyzed single glass grains from volcanic turbidites cored at ODP Site 840 for a series of major, trace and REEs (Fig. 1). The most primitive glass shards showed that HFSEs became increasingly depleted in the most incompatible elements between 7.0 and 4.0 Ma. This trend was reflected in the major element chemistry as a fall in the alkali element contents, interpreted to reflect variations in the degree of partial melting due to thinning of the arc lithosphere during Lau Basin rifting (cf., Plank and Langmuir, 1988). 4. NATURE OF THE VOLCANICLASTIC RECORD The sedimentary sequences that accumulate in forearc basins act as valuable records of the tectonic and magmatic evolution of the active plate boundary. Chemical analysis of redeposited volcaniclastic sediments has provided a new way of tracing arc chemical evolution (e.g., Hiscott and Gill, 1992) which can be readily correlated with tectonic events through dating of intercalated pelagic sediments. This approach hinges on the realization that the shards that comprise volcaniclastic sediments are samples of the volcanic glass erupted at the arc volcanic front at the time of deposition of the tephra or turbidite. Although not single-event tephra horizons, the turbidites incorporate shards from a modest number of eruptions that occurred over geologically short time spans. The redeposited character of the sediments at ODP Site 840 was discussed in detail by Clift (1994). A summary log is provided here to demonstrate large-scale variability appropriate to this study (Fig. 2A). The sands themselves comprise a mixture of volcanic chemistries, dominated by dacitic compositions, but with minor andesite, basaltic andesite, plagioclase and pyroxene grains (Clift et al., 1995). Texturally the grains are fresh (Fig. 2B) and show a vesicular fabric, most strongly in the high silica grains. Seismic data from Tonga (Tagudin and Scholl, 1994; Austin et al., 1989) suggest that large talus aprons form around the arc volcanoes, while further towards the trench the sediment sequences onlap the outer arc high from which substantial amounts of material may be episodically eroded (Clift et al., 1998). In practice sampling close to the arc centers, effectively within the debris aprons, ensures that the sediments represent the local volcanic output, with minimal “contamination” from along-strike or from eroded older sequences along the outer forearc high. ODP Site 840 is located close enough to Ata (⬃25 km) that it can be used as record of the activity of that center (Fig. 1). Unfortunately even at ODP Site 840 the volcaniclastic record of turbidites and debris flow conglomerates is not complete because of a hiatus from ⬃3.5 to 4.5 Ma (Fig. 2). Since rifting of the Lau Basin separated the original Miocene arc from the 3349 forearc basin a volcaniclastic hiatus developed in the forearc stratigraphy (e.g., Clift et al., 1995). Fortunately this gap in the record of the forearc can be filled by that preserved at ODP Site 834, which represents the oldest submarine volcanism within the Lau Basin, effectively the arc volcanic front, at a time when there was no subaerial arc. 5. BORON IN SUBDUCTION ZONES Boron has become an increasingly important element in tracing the influence of recycling material in island arc systems. Studies of fluids vented from subduction accretionary prisms indicate that most boron is held in an exchangeable form in the sedimentary cover or altered oceanic crust (AOC), but not in the mantle wedge (e.g., Spivack et al., 1987; Morris et al., 1990). Boron is lost mostly from the subducting plate at low temperatures (e.g., You et al., 1993, 1995), but a minor amount is structurally bound into minerals and can be transferred to the mantle. A complete mass balance requires that some exchangeable boron is subducted to relatively deep levels, especially in colder, faster subduction zones (Bebout et al., 1999). If subduction is rapid then boron dissolved in sediment pore-waters may be subducted further than the normal shallow regions below forearcs. Be is less fluid mobile than boron and is subducted to deeper levels, behaving much like a LREE (Tatsumi and Isoyama, 1988). Although partition coefficients of B and Be can differ by two orders of magnitude under certain conditions (Chaussidon and Libourel, 1993) the correlation of B/Be with other ratios indicative of slab-derived fluids (e.g., Ba/Ce, Rb/La) demonstrates that these elements have very similar mineral-melt partition coefficients in most subduction zone environments (Ryan, 1989; Ryan and Langmuir, 1993). Partial melting and fractional crystallization processes do not therefore significantly fractionate B from Be, and variations in the B/Be ratio in arc lavas are controlled primarily by differences in the slab input to the mantle sources of the lavas. Most arc lavas have significantly higher B/Be ratios than midocean ridge and oceanic island basalts, because boron is added to the source of the arc lavas by fluids derived from the subducting slab. The B/Be ratio is thus a useful indicator of the amount of slab-derived boron in arc lavas (Morris et al., 1990; Edwards et al., 1993; Gill et al., 1993; Hochstaedter et al., 1996). Correlations of the B/Be ratio with 10Be/9Be in some arc lavas (Morris et al., 1990; Leeman et al., 1994) suggest that boron is derived at least in part from subducted sediment, and is rapidly transferred from the subducting slab to the surface in lavas (within about 5 half-lives of 10Be, or 7.5 m.y.). Morris et al. (1990) note that modern oceanic sediments have B/Be ratios of 50 to 60, and occasionally close to ⬃100, so that values exceeding that level require input from an additional boron reservoir, probably the AOC. Boron isotopes have been used successfully to trace the source of fluids from the subducting plate in a number of arc systems (Ishikawa and Nakamura, 1994; Ishikawa and Tera, 1999). Source resolution is possible because of the large boron isotopic differences between the surface reservoir for boron (i.e., especially between deep sea sediments, the AOC and the mantle wedge). Although sediment ␦11B is variable, mostly depending on sediment composition (from –17.0‰ to ⫹4.5‰ 3350 Peter D. Clift et al. 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. Dark colors correspond to more mafic grains, although dacitic material dominates. for non-carbonate lithologies; Ishikawa and Nakamura, 1993), the mantle has relatively homogenous negative values of ␦11B (⫺10 ⫾ 2; Chaussidon and Marty, 1995). In contrast, AOC shows positive values of ␦11B (between 0.1–9.2 ⫾ 0.4‰; Spivack and Edmond, 1987; ranging up to ⫹24.9‰; Smith et al., 1995). These features make boron a valuable tracer of the origin of slab-derived material (melt and/or fluid) in subduction zones. Boron isotopic tracing of slab-flux in Tonga 5.1. Analytical Conditions The Cameca ims 3f at Woods Hole Oceanographic Institution was set-up for MRP⫽1800, energy slit, field aperture and contrast aperture settings following procedures developed by Chaussidon et al. (1997). The primary beam is accelerated to 10 keV, with a secondary accelerating voltage of 4.5 keV and produces a beam size of 20 to 40 m on the sample. Background interference levels are low, ⬃0.05 ppm, and matrix effects are negligible (Chaussidon et al., 1997). Replicate analyses of a silicate glass standard (GB4) have produced ␦11B ⫽ ⫺12.7 ⫾ 0.8‰ (2 standard error, n ⫽ 9), consistent with the accepted value of ␦11B ⫽ 12.80‰ (Chaussidon et al., 1997). The precision of individual analyses, involving 60 cycles of measurements (counting times were 15 s for 10B and 8 s for 11B for each cycle), ranges from ⫾0.6 per mil to ⫾2.2 per mil (2), comparable to those of Chaussidon et al. (1997). Variations in boron isotope ratios are described by the notation ␦11B where ␦11B ⫽ 冋 册 11 B/10B Sample ⫺ 1 ⫻ 1000 11 B/10B Standard (1) The standard used is NBS, the Searles Lake evaporite (McMullen et al., 1961). Results of the ion probe boron analysis are shown in Tables 1 and 2. In addition, subaerial volcanic rock specimens from the volcanic islands of the Tonga-Kermadec Arc were analyzed for a range of REEs and trace elements using the Cameca ims 3f microprobe. The results from these rocks are given in Table 2, with analytical uncertainties in Table 3. 6. RESULTS B/Be varies significantly along the strike of the arc, with a range of values seen at each volcano (Fig. 3). The Tonga and Kermadec segments of the arc have similar B/Be ratios around 15 to 95. B/Be never exceeds 122. Raoul in the Kermadecs shows 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 higher values than Tafahi, Tofua and Ata (14 –52, 28 –33, and 23–27 respectively). Temporal variation in B/Be measured in the volcaniclastic record (Fig. 4) shows that in fact modern volcanic B/Be values are small (⬍130) compared to earlier activity (values up to 520 at 3– 4 Ma). The record at ODP Site 840 shows a well developed increase in B/Be values from about 70 at 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, but highly variable, values (⬎500) are noted at the base of the section at ODP Site 834. Since 3 Ma there appears to be a gradual decline in B/Be values to the present day value of 27 at Ata, most notably at ODP Site 840. B/Be values measured in the youngest volcaniclastic layers are consistent with the value measured from modern Ata, the closest modern volcanic center (20 –30). Globally modern Tonga-Kermadec has a relatively high range of B/Be values (6 –122), although similar to those recorded from Nicaragua (Reagan et al., 1994), Papua New Guinea (Gill et al., 1993), and the Bismarck Arc (Morris et al., 1990). Although there is a temporal variation in B/Be at ODP Site 3351 840, the isotopic character of the volcanic glass does not show a systematic change with time through the drilled stratigraphy (Fig. 5). ␦11B values seen since 7 Ma at ODP Site 840 are mostly strongly positive (ranging from 0.6 –37‰), with occasional negative excursions at 1.7 Ma and 3.6 Ma (⫺1.3 and ⫺11.6, respectively). ␦11B values appear to be lower (from about ⫹1 to ⫹15‰) at the base of the section and then rise to higher values (⫹12 to ⫹25‰) by ⬃5.5 Ma. Since that time the values have been constantly positive and high, typically 10 to 30. Some single intervals show significant spread of isotopic values, e.g., ␦11B values from ⫺11.6 to ⫹32.8‰ in different grains in a single turbidite sand dated at 3.6 Ma. These analyses represent the first boron isotopic data from volcaniclastic deposits published. The ␦11B values are considerably more positive than existing boron isotopic data from island arc lavas (usually ⫺6 to ⫹7‰; Peacock and Hervig, 1999). ␦11B is dominantly positive along length of the modern arc (ranging from ⫺44‰ to 30.8‰), but is highest in central and southern Tonga (from ⫺4.4 to⫹30.8‰, Fig. 6A). Raoul in particular has lower ␦11B, but with a spread of values from ⫺4.4 to ⫹17.1‰. There is a good deal of scatter at each arc volcano, even within analyses conducted on the same sample. This trend may reflect either small scale magmatic variability or limited alteration. Rain water has ␦11B values of ⫹3.8 to ⫹9.1‰ (Rose, 1999), however its effectiveness as a contaminating agent in the case of the subaerial lavas is very much restricted 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 clearly concern regarding the meaning of boron analyses made in volcanic glass shards as small as 40 m across, which have spent significant amounts of time surrounded by seawater and diagenetic fluids. This concern is especially strong for silicic grains which tend to be more vesicular and thus have a high surface area/volume ratio. Several factors now suggest that the alteration of the glass grains considered here is low, at least in the upper part of the drilled section at ODP Site 840, and that the boron is representative of the magma source and not burial diagenesis. Visual inspection of the grains reveals clear pristine grains, unclouded by hydration, except in rare cases towards the base of ODP Site 840. However, chemical alteration of mobile elements may begin before visible change has occurred, and this may be seen as low analytical totals in electron probe analysis. In arc volcanic rocks low analytical totals can arise from an indigenous volatile content in the melt or from subsequent hydration. Sobolev and Chaussidon (1996) estimate that primary subduction melts contain 1 to 3% H2O, values that may be expected to rise during crystal fractionation (e.g., Burnham and Jahns, 1962), although much of this may be lost by degassing during eruption. In a varied suite of lavas evolving under different conditions of volatile degassing, and perhaps differing primary water contents, a range of volatile contents might be expected over much of the compositional spread, with a broad correlation between volatile-loss measured by [100%—Total], and silica content. Figure 7A shows this to be the case. The data has a sloping upper bound which corresponds to the effective 3352 Peter D. Clift et al. Table 1. B and Be concentration and isotopic data from SIMS analysis of single tephra particles from ODP Sites 834 and 840. 840A–1–1, 102cm 840A–1–1, 102cm 840A–1–1, 102cm 836–2–6, 38cm 840B–1–CC, 33cm 840A–1–2, 75cm 840A–1–2, 75cm 840A–1–2, 75cm 840A–1–2, 75cm 840A–1–3, 45cm 840A–1–3, 45cm 840A–1–3, 45cm 840B–1–CC, 33cm 840B–1–CC, 33cm 839–2–1, 44cm 836–3–6, 123cm 836–3–6, 123cm 838–3–2, 147cm 839–3–4, 22cm 835–4–2, 121cm 838A–5–5, 85cm 839–4–4, 35cm 834–3–1, 72cm 837–4–6, 67cm 835–8–6, 25cm 837–6–4, 132cm 839–10–3, 60cm 839–11–6, 82cm 835–14–6, 84cm 837–7–4, 33cm 838–7–2, 89cm 838–7–2, 89cm 840B–3X–CC, 10 cm 840B–3X–CC, 10 cm 840C–1H–2, 29cm 840C–1H–2, 29cm 840C–1H–2, 114cm 839–15–1, 79cm 835–16–3, 100cm 839–18–CC, 6cm 839–21–CC, 33cm 837–8–2, 49cm 838–11–5, 30cm 837–9–1, 53cm 837–9–3, 93cm 840C–2–1, 37cm 840C–2–1, 37cm 840C–2–1, 37cm 840C–3–3, 2cm 840C–3–3, 2cm 840C–3–3, 2cm 834–6–2, 35cm 840C–4–6, 34cm 835–15–6, 83cm 834–7–4, 67cm 840B–10X–CC, 4cm 834–9–3, 98cm 834–9–3, 98cm 834–9–4, 95cm 834–9–4, 95cm 834–9–4, 95cm 834–9–6, 92cm 834–9–6, 92cm 834–10–3, 102cm 834–10–3, 102cm 834–10–3, 102cm 834–10–CC, 13cm Age (Ma) Be (ppm) B (ppm) %2 0.35 0.35 0.35 0.40 0.40 0.43 0.43 0.43 0.43 0.45 0.45 0.45 0.45 0.45 0.50 0.60 0.60 0.75 0.75 0.80 0.80 0.95 1.00 1.00 1.45 1.60 1.68 1.70 1.70 1.70 1.70 1.70 1.70 1.70 1.70 1.70 1.73 1.73 1.73 1.74 1.74 1.90 1.95 2.00 2.05 2.20 2.20 2.20 2.70 2.70 2.70 2.90 3.00 3.05 3.30 3.40 3.50 3.50 3.52 3.53 3.53 3.53 3.55 3.57 3.57 3.57 3.58 0.198 0.198 0.198 0.012 0.020 0.018 0.018 0.018 0.018 0.021 0.021 0.021 0.020 0.020 0.025 0.019 0.017 0.025 0.026 0.025 0.022 0.028 0.020 0.017 0.024 0.023 0.016 0.023 0.037 0.022 0.016 0.024 — — 0.013 0.013 — 0.014 0.028 0.017 0.030 0.018 0.017 0.015 0.017 0.012 0.012 0.012 0.011 0.011 0.011 0.021 — 0.035 0.033 0.013 0.029 0.029 0.018 0.018 0.018 0.018 0.018 0.022 0.022 0.022 0.016 4.70 4.70 4.70 2.31 2.59 2.17 2.17 2.17 2.17 2.57 2.57 2.57 2.59 2.59 1.19 2.19 1.31 0.79 0.97 1.18 1.25 1.64 3.69 3.72 2.98 0.77 3.69 0.59 5.04 1.30 0.62 1.96 — — 3.14 3.14 — 0.91 1.40 0.98 0.88 0.73 5.32 4.01 1.19 3.87 3.87 3.87 4.79 4.79 4.79 3.15 — 5.59 3.44 5.48 7.27 7.27 4.91 4.91 4.91 8.75 8.75 3.73 3.73 3.73 3.39 0.39 0.78 — — — — 0.47 0.70 — 0.73 — — — 0.41 — — — — — — — — — — — — — — — — — — 0.83 0.88 0.33 — 0.66 — — — — — — — — — 0.29 0.37 0.55 — — — 0.61 — — 0.23 0.22 — — 0.36 0.76 0.45 0.63 0.55 0.53 — — 11 B/10B ␦11Boron 3.911 3.906 — — — — 3.963 3.936 9.097 6.814 — — — — 22.497 14.490 3.979 — — — 3.937 — — — — — — — — — — — — — — — — — — 3.946 3.928 3.966 — 3.917 — — — — — — — — — 3.990 3.991 3.960 — — — 3.967 — — 3.967 3.922 — — 3.988 3.966 3.976 3.983 4.003 3.882 — — 26.634 — — — 15.742 — — — — — — — — — — — — — — — — — — 5.869 1.468 23.210 — ⫺1.329 — — — — — — — — — 29.439 15.940 21.650 — — — 23.541 — — 23.541 12.019 — — 28.863 23.163 18.292 20.046 32.853 ⫺11.633 — — Volatile % 2.49 6.76 8.30 3.05 3.37 0.10 15.34 0.40 11.17 0.22 2.95 1.71 0.86 2.66 7.90 4.39 5.75 6.48 5.34 7.16 5.96 6.80 2.89 3.76 5.21 2.33 5.65 3.95 4.79 3.10 4.82 3.76 4.93 7.50 4.93 7.50 3.17 3.01 6.33 3.67 3.82 1.83 5.39 4.24 2.55 5.22 5.65 5.76 4.18 2.31 4.92 6.28 1.09 7.24 3.09 6.29 7.86 6.74 6.16 9.50 7.83 7.58 7.58 5.19 6.90 7.84 4.10 (Continued) Boron isotopic tracing of slab-flux in Tonga 3353 Table 1. (Continued) 834-10-CC, 13cm 834–10–CC, 13cm 834–1o–CC, 13cm 834–11–1, 40cm 834–12–2, 88cm 834–12–2, 88cm 834–12–2, 88cm 840B–11–1, 122cm 840B–12–5, 49cm 840B–12–5, 49cm 840B–12–5, 50cm 840B–13–1, 137cm 840B–13–3, 137cm 840B–13–3, 137cm 840B–13–3, 137cm 840C–5–CC, 7cm 840C–5–CC, 7cm 840C–6–4, 123cm 840C–7–4, 128cm 840B–17X–CC, 2cm 840B–17X–CC, 2cm 840C–8–1, 97cm 840C–9–1, 2cm 840C–10–1, 3cm 840C–12–2, 132cm 840B–26–1, 60cm 840C–13–5, 79cm 840B–28–1, 77cm 840C–13–5, 79cm 840B–30–1, 12cm 840B–33–1, 21cm 840B–33–1, 21cm 840B–34X–1, 24cm 840B–37X–1, 13cm 840B–39–1, 53cm 840B–39–1, 53cm 840B–46X–2, 9cm 840B–46X–2, 9cm 840B–52X–2, 5cm 840B–52X–2, 5cm 840B–52X–2, 5cm Age (Ma) Be (ppm) B (ppm) %2 3.58 3.58 3.58 3.70 3.80 3.80 3.80 4.50 5.00 5.00 5.00 5.05 5.05 5.05 5.05 5.08 5.08 5.09 5.11 5.13 5.13 5.13 5.25 5.30 5.33 5.37 5.38 5.38 5.38 5.40 5.65 5.65 5.80 5.90 5.98 5.98 6.30 6.30 6.53 6.53 6.53 0.016 0.016 0.016 0.018 0.014 0.014 0.014 0.020 0.022 0.023 0.026 0.023 0.023 0.023 0.023 0.022 0.022 0.025 0.018 0.034 0.034 0.020 0.013 0.014 0.019 0.019 0.014 0.018 0.014 0.013 0.019 0.034 0.025 0.043 0.020 0.023 0.041 0.030 0.023 0.023 0.035 3.39 3.39 3.39 10.84 7.36 7.36 7.36 4.64 2.08 0.41 2.09 4.04 4.04 4.04 4.04 3.39 3.39 4.20 4.59 4.74 4.74 4.33 2.37 2.39 3.13 3.80 3.15 6.35 3.15 1.85 2.13 1.46 0.64 1.62 2.50 2.89 3.48 2.03 1.89 1.89 0.91 0.38 0.41 — — 0.39 upper limit of SiO2 in a granitic liquid when subject to a variable addition of volatiles. This plot allows very anomalous low totals at a particular silica content to be identified. This plot does not sort bad from good analyses but instead only identifies those grains that are almost certainly significantly hydrated during diagenesis. Although diagenesis may affect grains throughout the section, the temporal distribution of those with anomalously low totals shows a clear pattern of strongest alteration of the oldest grains, as might be expected (Fig. 7B). There is no correlation between high B/Be and the zone at the base of the section at ODP Site 840 where alteration is most common. 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 content and boron isotopic ratios (Fig. 7C). However, there is some correlation between lower ␦11B and higher alteration at the base of ODP Site 840 (Fig. 7D), consistent with the observation that diagenesis and alteration of volcanic glass tends to decrease ␦11B values (Ishikawa and Nakamura, 1993). In Figure 7D the amount of excess volatiles beyond the maximum that can reasonably be associated with fractional crystallization is ␦11Boron Volatile % 3.913 3.932 — — 3.962 9.535 14.420 — — 22.275 — 0.23 1.00 — — — 0.29 — 3.999 3.964 — — — 3.989 — 31.802 9.157 — — — 29.101 0.60 0.58 — — — — — 0.38 — — — — 0.41 0.59 0.59 0.88 — — — — — 0.71 — 0.32 0.32 0.80 3.957 4.022 — — — — — 3.991 — — — — 4.026 3.973 3.923 3.905 — — — — — 3.960 — 3.942 3.930 3.917 20.800 37.576 — — — — — 29.758 — — — — 24.877 17.485 12.192 7.475 — — — — — 8.125 — 17.097 0.604 3.262 7.66 6.24 3.65 6.00 9.29 6.52 6.22 8.10 3.92 2.81 2.98 6.48 3.70 6.48 3.70 4.53 6.71 4.74 2.71 2.71 0.15 0.00 3.10 3.37 6.56 6.49 6.49 6.17 5.19 2.96 5.89 5.09 5.80 3.22 12.10 2.06 0.62 8.39 5.51 5.98 4.71 11 B/10B compared with ␦11B. The least altered grains (i.e., with high negative excess volatiles, ⬍⫺5%) show a range of ␦11B values between 20 and 29‰. However, the lowest ␦11B values (⬍6‰) are found only in grains with the higher volatile excess (⬎⫺4% excess volatile), i.e., the most altered grains tend to have lower ␦11B values. The higher values that dominate the upper part of the hole can be considered to be the most magmatic. Volcaniclastic ␦11B values in sediments younger than 5.5 Ma, typically with the lowest excess volatile content, are close to those ␦11B values measured from the subaerial volcanic rocks of adjacent Ata (8.6 and 16.4‰), consistent with their composition being close to magmatic values. If the ␦11B in volcanic sediments younger than ⬃5.5 Ma had been affected by diagenesis then these might be expected to have fractionated to ␦11B values more negative than those we can be reasonably certain are primary and magmatic. We conclude that for sediments younger than 5.5 Ma ␦11B may be considered to be close to 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 in the Marianas (Clift and Lee, 1998). This peak may be a re- Sample # T114 T114 T114 Fon–31 Fon–31 Fon–31 Fon–31 Fon–8–69 Fon–8–69 Fon–8–69 Tofua–32 Tofua–32 Tofua–32 HHTop HHTop HHTop HHTop HHTop 582–8–4 582–8–4 582–8–4 582–8–4 — — — 23374 23374 23374 23374 23383 23383 23383 14837 14837 14837 Island Tafahi Tafahi Tafahi Fonualei Fonualei Fonualei Fonualei Fonualei Fonualei Fonualei Tofua Tofua Tofua H. Ha’apai H. Ha’apai H. Ha’apai H. Ha’apai H. Ha’apai Ata Ata Ata Ata Seamount Seamount Seamount Raoul Raoul Raoul Raoul Raoul Raoul Raoul L’Esperance L’Esperance L’Esperance 1.05 1.52 0.65 6.49 5.55 6.61 5.63 7.60 1.87 5.43 5.41 3.21 2.68 7.81 1.26 10.37 6.52 10.97 9.36 6.81 6.20 9.50 3.62 4.98 3.52 6.64 5.01 11.24 4.50 5.30 0.88 1.19 2.00 1.24 0.96 3.15 0.50 3.91 2.58 3.97 3.89 2.60 2.15 3.94 1.12 1.50 1.06 2.18 1.79 4.69 1.15 2.50 Ce 0.44 0.73 0.38 2.78 2.38 2.87 2.11 3.08 La 6.88 2.64 3.23 6.52 4.67 9.09 5.75 5.15 4.17 6.05 3.07 4.79 3.65 5.12 3.53 1.19 4.21 4.38 2.23 2.65 5.20 0.92 1.11 1.05 0.70 4.14 3.29 4.35 4.31 4.68 Nd 2.50 1.20 1.18 2.19 1.80 3.10 1.38 1.73 1.27 1.73 1.38 2.14 1.40 2.02 1.20 0.28 1.60 1.65 0.69 1.41 1.64 0.57 0.54 0.39 0.38 1.35 1.06 1.23 1.43 1.56 Sm 2.06 0.69 2.47 2.94 1.44 2.61 1.54 2.27 1.84 1.48 1.08 1.63 0.84 1.98 0.78 0.61 1.59 0.83 0.94 1.01 1.77 0.49 0.53 0.54 0.33 2.08 1.44 1.18 2.79 2.82 Eu 3.82 2.30 1.45 3.00 2.53 6.01 2.00 3.06 1.78 2.72 2.83 4.11 2.93 3.45 1.59 0.35 2.32 2.25 1.24 2.67 2.69 0.72 1.34 0.76 0.77 2.02 1.45 1.82 1.97 1.97 Dy 2.36 1.82 1.09 1.78 1.59 4.19 1.50 2.11 1.29 2.10 1.86 2.52 2.22 2.22 1.22 0.22 1.42 1.53 1.02 1.92 1.91 0.55 0.90 0.56 0.62 1.19 0.94 1.53 1.38 1.24 Er 2.52 2.42 0.88 2.12 1.96 4.83 1.44 2.46 1.49 2.28 2.16 2.93 2.29 2.62 1.32 0.18 1.71 1.62 0.95 2.12 1.95 0.78 1.21 0.73 0.74 1.72 1.05 1.57 1.70 1.39 Yb 0.1458 0.1574 0.0578 1.4069 0.1027 0.1000 0.0849 0.0971 0.1018 0.1274 4.06 10.39 8.51 2.99 5.21 2.71 4.73 5.26 0.03 0.06 0.1325 0.1154 0.0794 0.0425 0.0359 0.16 0.29 0.04 0.87 4.62 0.1642 0.1258 0.1114 0.1048 0.1073 0.0726 0.0844 0.0922 0.0161 0.0958 0.0871 0.1134 Be 1.20 0.25 1.07 6.71 3.99 3.15 6.92 11.10 3.17 8.99 1.90 1.75 Li 8.021 3.418 3.183 8.166 3.938 12.215 8.151 8.214 2.562 3.729 3.0475 3.1233 4.342 3.249 1.983 8.649 2.183 1.644 6.725 4.126 5.014 6.841 8.000 0.481 3.270 2.516 3.835 B 1394.3 603.3 5421.3 2787.2 2838.3 1282.4 1399.8 239.1 15720 2519.5 1897.6 1204.0 864.5 850.5 394.4 341.0 618.4 1221.6 1069.6 Ti 1.8 1.1 2.8 16.3 7.9 5.7 5.3 4.1 2.7 2.7 0.9 1.7 0.9 0.8 0.9 0.7 6.5 3.8 6.5 6.4 8.6 4.4 4.5 3.1 3.8 Rb 155.3 147.7 50.1 90.0 73.9 183.5 109.5 80.8 124.4 132.1 129.0 46.6 45.0 139.0 127.1 103.7 187.1 159.7 76.6 90.8 91.1 25.5 34.5 141.7 121.5 Sr 11.9 4.6 6.0 7.9 9.1 79.4 17.2 17.4 6.2 7.0 1.3 2.4 1.2 1.7 2.3 2.0 6.6 8.3 5.1 12.8 16.4 12.5 3.2 8.2 11.9 Y 68.8 5.8 12.3 27.3 16.7 73.4 49.6 36.3 14.7 20.3 1.1 10.3 0.8 3.5 4.6 3.7 21.4 17.2 13.6 33.1 41.4 4.7 9.3 19.6 21.9 Zr 0.215 0.039 0.300 8.128 1.160 0.541 0.404 0.254 0.111 0.241 0.026 0.385 0.117 0.215 0.137 0.117 0.372 0.241 0.365 0.567 0.671 0.033 0.795 0.202 0.176 Nb 171.0 40.9 48.6 105.4 74.7 139.3 130.9 69.9 41.5 49.4 28.5 21.4 11.5 21.7 14.1 11.7 148.8 108.6 54.4 172.8 230.5 11.1 27.0 89.8 93.0 Ba 55.00 21.72 55.04 5.80 38.36 122.16 96.02 84.61 25.17 29.27 23.13 27.06 54.67 76.44 55.23 52.68 17.35 14.76 64.17 38.47 69.08 81.02 86.81 29.83 34.11 28.87 33.81 B/Be Table 2. Trace element, REE and B and Be concentration and isotopic data from SIMS analysis of volcanic rocks from the Tonga-Kermadec Arc. 0.45 0.46 0.31 0.42 0.45 0.23 0.47 0.39 0.36 0.36 0.26 0.82 0.48 0.51 0.72 0.71 0.51 0.69 0.23 0.12 0.42 0.52 0.71 0.64 0.68 0.83 %2s 3.9865 3.9729 3.9185 3.9422 3.9427 3.9494 3.8849 3.8857 3.8983 3.9055 3.9204 3.9336 3.9382 3.9427 3.9952 3.9404 3.9726 3.9657 3.9843 3.9692 3.9642 3.9610 3.9396 3.9380 3.9325 3.9538 3.9554 3.9479 3.9497 3.9672 B/10B 11 28.53 24.03 10.99 17.11 16.25 6.83 2.33 2.54 4.80 ⫺4.36 11.48 14.89 16.06 17.23 30.77 16.64 23.96 23.16 27.97 23.08 10.60 14.49 16.43 8.61 14.61 19.10 13.05 17.60 11.61 16.07 ␦11B 3354 Peter D. Clift et al. Boron isotopic tracing of slab-flux in Tonga 3355 Table 3. (Analyses from 135– 840A–1H–3, 45cm:Hf from 135– 840B–35X–3, 22 cm) Element Rb Sr Y Zr Nb Ba La Ce Nd Sm Eu Gd Hf Tb Dy Ho Er Tm Yb Lu Pb Th U Typical % error in count rates Concentration (ppm) 2.4 0.4 1.0 0.9 4.7 1.0 3.3 2.3 7.5 12.1 8.6 9.3 8.5 9.3 10.0 9.1 11.5 13.7 11.5 15.8 2.6 26.7 24.3 7.040 214.0 14.40 60.00 0.576 188.0 3.920 9.780 6.100 1.670 0.112 2.900 1.990 0.406 1.800 0.522 1.550 0.245 1.560 0.181 1.160 0.174 0.226 gional feature of active margin volcanism at this time, or may be linked to rifting of the respective backarc basins. If the 3 to 4 Ma peak B/Be seen at ODP Site 840 were related to alteration, it would be remarkable that the same effect is seen not only at ODP Site 834, but also in a different arc and different borehole at different burial depths, but at the same stage in the arc’s evolution. We therefore suggest that the B/Be record of the tephra is not significantly damaged by the diagenesis, at Fig. 4. Temporal variation in B/Be in the Tonga system as sampled at ODP Site 840. The record shows peak values at 3 to 4 Ma, just after break-up of the Lau Ridge Arc. Values at the volcanic front have decreased since 3 Ma to the modern values. least back to ⬃5.5 Ma. The super-high B/Be values noted (up to 607) must therefore represent a phase of extreme slab flux into the arc, which does not appear to have an equivalent in the modern oceans. Fig. 3. Along-strike variations in B/Be in the Tonga-Kermadec Arc measured by SIMS analysis of volcanic glass and fine grained groundmass of modern arc volcanic rocks. 3356 Peter D. Clift et al. Fig. 6. Along-strike variations in (A) boron isotope ratios in the Tonga-Kermadec Arc, as measured by SIMS analysis of volcanic glass and fine grained groundmass of modern arc volcanic rocks. Variations in (B) 230Th/232Th and (C) 87Sr/86Sr for the Tonga Arc. Data from Regelous et al. (1997) and Ewart et al. (1998). Fig. 5. Temporal variation in boron isotope ratios in the Tonga system, showing a generally constant, positive ␦11B ratio since 7 Ma. 7.2. Volume of Fluid Flux The high B/Be values in Tonga-Kermadec lavas are interpreted to represent generally high volumes of modern slab fluid-flux, with up to 95% of the total boron being slab-derived fluids. In contrast, B/Be values for the depleted upper mantle are ⬍5 (Hochstaedter et al., 1996). The relative volume of slab fluid-flux charted by B/Be at ODP Site 840 shows long term variability. The very high B/Be values of the sediments deposited at 3 to 4 Ma are about three times higher than reported for any other arc lavas yet analyzed, suggesting temporary dramatic slab fluid-flux at that time (B from the slab comprising 99% of total). They are also far in excess of known sediment values which are typically 50 to 60, and reach a maximum of ⬃100 (Morris et al., 1990), requiring an additional source for the excess boron. In contrast, known B/Be values from AOC range from 100 to 400 (Fig. 7; Thompson and Melson, 1970; Seyfried et al., 1984). B/Be values in volcanic glass are low in slow subduction zones, e.g., Aeolian, Campanian and Aegean Arcs (Morris et al., 1993, Clift and Blusztajn, 1999), or where the subducting plate 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., Marianas (Clift and Lee, 1998), Bismarck Arc (Morris et al., 1990). Although B/Be values are affected by several factors it appears that there is a first order correlation between B/Be and the thermal structure of the subduction zone, a consequence of the temperature sensitivity of boron in subducting sediments. Experimental data indicate that at elevated temperatures (200 – 350°C) most boron in subducting sediment has been mobilized (You et al., 1995), while more metamorphosed sediments showing progressive loss, so that by ⬃800 to 900°C the vast majority of exchangeable boron in sediments or altered volcanic rocks has been lost (Moran et al., 1992). By the time the slab reaches its melting point little excess boron remains. In these settings boron is subducted to greater depths, where it may be introduced into the arc lava source. The Tonga-Kermadec arc is likely to have a low thermal gradient because it is subducting relatively old and cold oceanic lithosphere and because the rate of convergence, especially in the north, is very fast (⬃20 cm/yr.; Pelletier et al., 1998). The high B/Be in the modern Tonga-Kermadec arc, as well as in the volcaniclastic record, argues strongly for slab dewatering and against slab melting under the arc volcanic front. Boron isotopic tracing of slab-flux in Tonga 3357 Fig. 7. Diagram showing relationship between ␦11B (origin of slab flux) and B/Be (volume of slab flux). High B/Be is always associated with positive ␦11B and a dominance of water derived from the altered oceanic crust. Black dots represent analyses from ODP Site 840 volcaniclastic sediment, circles represent modern arc lavas. Black squares represent end-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 crust data from Thompson and Melson (1970) and Seyfried et al. (1984). 7.3. Fluid Sources The isotopic boron compositions and B/Be values in volcanic glass can be used as a guide to the source of the slab flux, because the mantle wedge, subducted sediments and AOC have very different boron concentrations and isotope compositions. These signals are not transmitted directly to the surface because boron isotopic fractionation at low temperature produces fluids derived from the slab with higher ␦11B than their source (e.g., Peacock and Hervig, 1997, 1999). Those fluids could trigger the melting of the mantle wedge and result in high ␦11B in arc lavas. The degree of isotopic fractionation of the fluid expelled from the slab is a function of the temperature (Oi et al., 1989), with the greatest fractionation occurring at low temperatures. The very high ␦11B seen along-strike is consistent with the cold state of the Tonga-Kermadec subduction zone. Isotopic fractionation does not occur when melting occurs because there is no difference in compatible in siliceous melts. The very high ␦11B seen in this study are higher than any possible source and require transfer to the mantle wedge in aqueous solution, not by direct melting of the subducting material. Measured ␦11B is rather higher than other known oceanic arc volcanic compositions (e.g., Izu, ␦11B ⫽ ⫹1.2 to ⫹7.3‰; Ishikawa and Nakamura, 1994; Halmahera, ␦11B ⫽ ⫺2 to ⫹4‰; Palmer, 1991; Martinique, ␦11B ⫽ ⫺6 to ⫹2‰; Smith 3358 Peter D. Clift et al. et al., 1997; Kurile, ␦11B ⫽ ⫺4 to ⫹6‰; Ishikawa and Tera, 1997). However, at low temperature (25°C) Peacock and Hervig (1999) have demonstrated that fluids with ␦11B as high as ⫹20‰ can be expelled from sediments with initial ␦11B of 0‰, even more positive than the ⫹15‰ ␦11B values measured in natural samples by Spivack et al. (1987) for the exchangeable boron released from sediments in accretionary complexes. The boron isotopic composition of AOC, with initial ␦11B typically between ⫺5 and ⫹10‰, can also undergo isotopic fractionation during dehydration. The fluid expelled will have more positive ␦11B value than the subducting slab (following a Rayleigh distillation model, e.g., Peacock and Hervig, 1999). Although temperatures under the arc volcanic front might be expected to be higher than those that yield the most extreme isotopic fractionation, high ␦11B values can be incorporated into arc volcanic rocks if the base of the hydrated mantle wedge is dragged down with the subducting slab, as suggested by Tatsumi (1989). This is more likely in faster, colder subduction systems. Convergence at Tonga-Kermadec (20 cm/yr.) is significantly faster than Izu (8 –10 cm/yr.), despite the similar thermal age of the oceanic lithosphere in each case, making this arc most likely candidate for the most positive ␦11B worldwide. Positive ␦11B are reported in carbonates (Hemming and Hanson, 1992; Gaillardet and Allègre, 1995), but their absence from the stratigraphy of the Pacific abyssal seafloor offshore Tonga (Burns et al., 1973) makes carbonates an unlikely candidate for the source rock of the lavas with high ␦11B values. The pattern of positive ␦11B along the modern Tonga arc most likely reflects a fluid-flux dominated by the dehydration of the AOC similar to what has been proposed for the Mariana Arc (Ishikawa and Nakamura, 1994, 1999), but with a greater contribution of sediment dehydration in the southern extreme where ␦11B is lower. Li isotope work in the Izu Arc supports the hypothesis of fluid-flux from the AOC as being dominant in that oceanic arc (Moriguti and Nakamura, 1998). Unfortunately we do not have boron isotopic data from the abyssal seafloor immediately east of the trench. Instead we refer to the pelagic sediment data of Ishikawa and Nakamura (1993), made in the central Pacific. The wind-blown origin of deep sea clays makes them particular uniform in composition over wide areas, so that no significant error is expected as a result of this approximation. Modern marine sediment ␦11B values range from ⫺6.6‰ (clay) to ⫹10.5‰ (carbonate), although pelagic clays have uniformly negative ␦11B values (Ishikawa and Nakamura, 1993). The high positive ␦11B values of the arc lavas contrast with the low ␦11B values of the clays and argues against them being a major contributor to the boron budget of the arc. Without carbonate sediments the only appropriate boron reservoir is the fluid expelled during the dehydration of the AOC. Input of continental material eroded from New Zealand forms a plausible explanation for the more negative ␦11B values measured in the Kermadec Arc. The Kermadec Trench is largely filled with sediment, especially at its southern end, in contrast to the relatively empty Tonga Trench, in which even the igneous basement to the subducting plate is exposed in flexural fault scarps (Hawkins et al., 1999). In addition, because of the slower extension in the Kermadec backarc (Havre Trough) compared to Tonga the subduction zone might be expected to have a warmer thermal gradient, so that the boron isotopes are less susceptible to isotopic fractionation. Together these factors drive the Kermadec lavas to lower ␦11B values. 7.4. Mixing End-Member Compositions Figure 6 shows the relationship between ␦11B and B/Be. Although there is a lot of scatter at low values of B/Be, it is clear that all samples with high B/Be values also have high positive ␦11B, and that all samples with negative ␦11B have relatively low B/Be (⬍200). This trend is not due to seawater contamination. A simple mixing model between seawater (B ⫽ 4.5 ppm, ␦11B ⫽ 40‰, and B/Be ⫽ 16200) and an arbitrary end-member representing a melt produced under a low slab fluid-flux with B ⫽ 1 ppm, ␦11B ⫽ 0‰ and B/Be ⫽ 25, suggests that a contamination of 0.4% of seawater would increase drastically the B/Be (from 25–310), covering much of the measured B/Be variation. However, it would only increase the ␦11B value from 0 to ⫹0.7‰; this hardly represents 2% of the total measured ␦11B variation. Similarly, the boron concentration would increase by less than 2%. The trend on Figure 6 has a much steeper slope than the slope that seawater contamination would produce. In contrast, if we mix the arbitrary end-member with AOC (B ⫽ 13 ppm, ␦11B ⫽ ⫹8‰, Spivack and Edmond, 1987, and B/Be ⫽ 350), it appears that a fraction between 1 and 45% of AOC can explain most of the B isotopic and the B/Be variations. Nonetheless, mixing the AOC directly would produce ␦11B values slightly lighter than the ones measured in the tephra (Fig. 6). If AOC is dehydrated at the shallow levels of the subduction zone where the temperature is low, the fluid expelled will have a heavier ␦11B than the initial AOC due to isotope fractionation (Palmer et al., 1987; Oi et al., 1989). Depending on the extent of dehydration and the relative partition coefficient of B and Be between AOC and fluid phase, it possible to produce the measured ␦11B values (Fig. 6). Although we can not totally rule out the contribution of fluids expelled by dewatering of sediments, which could also produce a fluid with higher ␦11B than the initial sediments, the B isotopic fractionation required to reach the high ␦11B observed would have to be more extreme. This implies that to satisfy the boron data from the tephra, the model requires preferential involvement of the altered fraction of the subducting AOC, through derivation of a fluid. In summary, the relatively high ␦11B of Tonga-Kermadec lavas, compared to other arc lavas, results from the high rate of plate convergence and the Cretaceous age of the oceanic lithosphere that result in a cold subduction zone that accentuates isotopic fractionation. High ␦11B also reflects the lack of a significant thickness of isotopically negative, continentallyderived sediments along much of the Tonga Trench. 7.5. Influence of Louisville Ridge Subduction and southward migration of the Louisville Ridge along the Tonga forearc (Herzer and Exon, 1985) might have been expected to increase sediment subduction due to the presence of its sediment apron and as a result of the tectonic erosion and subduction of volcaniclastic sequences within the forearc. Louisville Ridge is a Cretaceous-Late Cenozoic hotspot related feature (Lonsdale, 1988) with more than 60 vol- Boron isotopic tracing of slab-flux in Tonga canoes arranged at intervals ⬍100 km apart along a 75-kmwide band stretching ⬃4300 km towards the SE from Tonga. The Tonga Trench is in collision with Osborne Seamount just south of Ata (Lonsdale, 1986). On the scale of the study presented here Louisville Ridge may be considered as a point source of material flux to the trench. Packham (1985) estimated that the Louisville Ridge began to collide with the northernmost part of Tonga forearc shortly after 4 Ma, and swept rapidly down the margin to its present location. Nonetheless, there is no low ␦11B measured at Tafahi, as might be expected for the involvement of larger volume of sediment with negative boron isotopic composition in petrogenesis. The Pb and Sr isotopic character of Tafahi and Niuatoputapu volcanic rocks has been used to argue that Louisville Ridge AOC or volcaniclastic sediment was influential over petrogenesis in this area (Ewart et al., 1998; Regelous et al., 1997; Turner and Hawkesworth, 1997). Figures 8B and C show the along-strike variation in 230Th/232Th and 87Sr/86Sr, as measured by Regelous et al. (1997) and Ewart et al. (1998). The increasing87Sr/86Sr and decreasing 230Th/232Th correlate with falling ␦11B in the Kermadec Arc and support a relationship between the proportion of sediment in the slab flux and ␦11B there. Comparison with Indian and Pacific mantle values (87Sr/ 86 Sr ⬍0.7034) suggests that the Sr system in the Kermadec Arc is strongly contaminated by sediment, whereas decreasing 87Sr/ 86 Sr moving northward along the Tonga Arc can be related to contamination by the lavas and volcaniclastic sediment of the plume-derived Louisville Ridge. However, at the north end there is no correlation in the Tonga Arc between the Sr and Th, and ␦11B. Our result suggests that the boron isotopic characteristics of the Louisville volcaniclastic apron differ from the continental sediments in the Kermadec Trench and are similar to the AOC of the Louisville Ridge itself. Given that the ultimate source of the boron in the Pacific AOC and the Louisville Ridge AOC is the same, i.e., seawater, it is not surprising that dewatering of Louisville Ridge does not affect the boron isotopic character of the volcanic rocks at Tafahi and Niuatoputapu. In contrast, Pacific AOC and Louisville AOC have totally different Sr, Th and Pb isotope compositions (Regelous et al., 1997), so that the Sr, Th and Pb isotope systematics of these islands are disrupted by ridge subduction. 7.6. Temporal Variations in Fluid Source Typically high ␦11B values at ODP Site 840 suggest that slab flux has been dominated by water expelled from the subducting AOC at least in the region of Ata since 7 Ma. Brief excursions to low ␦11B values are noted, with some sediments showing huge internal variations in ␦11B values. These sediments incorporate material from several eruption events, possibly incorporating material transported far along the arc or reworked from older deposits, since the degree of variability would be hard to account for in a single eruption. Short-term variations in ␦11B are likely due to changes in the proportion of sediment and AOC-derived fluids. This is because the lag time required to heat or cool a subduction zone following changes in the age of the subducting slab and its rate of convergence is too long to affect the boron isotope character of the arc over short time scales (Peacock, 1996). Given the composition of the modern 3359 Tonga Arc, the large distance of ODP Site 840 from New Zealand (and thus probable lack of input from sediments derived from there) and the apparent inability of volcaniclastic sediment to strongly alter the ␦11B values of the arc, this temporal history is not surprising. The relative lack of shortterm variations in ␦11B contrasts with the conclusions of Ishikawa and Nakamura (1999) in the Marianas Arc. We suggest that the relatively featureless character of the subducting plate in Tonga, compared to the seamount strewn Pacific east of the Marianas, results in a more constant supply of water dominated by the AOC. 8. INFLUENCE OF FLUID FLUX ON PETROGENESIS We now examine the coherence of variations in the slab fluid-flux with other elemental groups to determine the effect that fluid-flux has on the arc chemistry. Slab fluid-flux to the backarc is generally significantly less than that to the main arc volcanic front (Pearce et al., 1995). Consequently the volcanic front is the best place to examine the influence of fluid flux, as charted by boron, over other element groups. Clift and Dixon (1994) noted an up-section trend to decreasing alkali element concentrations from a maximum at 7 Ma, to a minimum at 3 to 4 Ma. K and Na show good negative correlation with B/Be, i.e., the volume of slab fluid-flux (Fig. 9. Contrary to their behavior in many arcs the alkali elements behave dominantly as incompatible elements and do not show enrichment by fluid-flux from the slab. Instead the falling maximum values reflect either increasing degrees of depletion in the mantle wedge or increasing degrees of melting of a relatively constant mantle source. Increased fluid flux can account for the trend by increasing partial melting through lowering of the solidus at any given depth in the wedge. The slab-derived fluids cannot have contained large amounts of alkali elements, despite their mobility in 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 ocean ridge basalts (Melson et al., 1976). The low concentration of the alkali elements in the arc lavas is thus consistent with the majority of the slab fluid-flux being derived from the AOC. To minimize the effects of fractional crystallization on the image of mantle melting conditions provided by the arc lavas and tephras we chose to examine the relative enrichment of HFSEs and REEs in only the more basaltic tephra shards (SiO2 ⬍60%) from the ODP wells. Figure 10 shows the relationship between Nb/Zr and La/Sm, and B/Be. Nb/Zr and La/Sm may be considered proxies for the relative enrichment in the HFSEs and REEs respectively. Both groups show highest depletion when B/Be is high (i.e., at 3– 4 Ma), although there is significant scatter. Such a relationship is in accord with the alkali element data (Fig. 9) in suggesting higher degrees of partial melting when slab fluid-flux is high. Moreover, the positive correlation of B/Be and both Nb/Zr and La/Sm is consistent with the measured B/Be being magmatic and not substantially affected by diagenesis. We propose an important modification to models that link REEs and slab flux, notably those in which LREE enrichment is dominated by sediment dewatering or melting (e.g., Plank et al., 1993; Hole et al., 1984). Instead our new data demonstrate LREE depletion driven by dewatering mostly of the AOC, as shown by high B/Be and ␦11B values in the volcaniclastic 3360 Peter D. Clift et al. Fig. 8. (A) Diagram showing the variations in silica versus the percentage of volatiles at ODP Site 840, estimated by the 100%-total measured major element composition. A positive correlation of SiO2 and volatiles is a natural progression of the fractionation process. Those grains with anomalously high volatile content, due to alteration can then be identified. (B) The temporal variation in volatiles shows that the vast majority of the altered grains were deposited before 5 Ma. (C) No correlation is noted between boron isotopic ratios and degree of volatile content. (D) Some of the least altered grains show very high ␦11B values suggesting that these are original and magmatic values. Boron isotopic tracing of slab-flux in Tonga 3361 Fig. 10. Diagram showing the temporal variations in total K2O contents for glass shards of all compositional ranges at ODP Site 834 and 840. Note the minimum values at 3 to 4 Ma. Data from Clift and Dixon (1994). 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-grain basaltic andesite shards.- shards. The very positive ␦11B and high B/Be values measured argue against slab melting in a hot subduction, because that would prevent the B isotope fractionation required here (Peacock and Hervig, 1999). In areas where slab melting is known to be important (e.g., western Aleutians, Yogodzinski and Kelemen, 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 correlate with the peak time of sediment subduction (⬃2–3 Ma) noted by Clift and Vroon (1996). Even if the temporal correlation is accepted, peak sediment flux might be expected to cause LREE enrichment, 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 volume of fluid fluxed from the subducting slab, and tracked by B/Be, can explain the chemical variability by controlling the degree of partial melting, since addition of water to the source mantle will depress the solidus. In practice this is the fluid-flux melting hypothesis of Luhr (1992), Ryan and Langmuir (1993) and Stolper and Newman (1994). Explanations of the chemical variability based on the rifting tectonics of the Lau Basin controlling the height of the melting column (e.g., Clift and Dixon, 1994) are consistent with the LREE and HFSE data but do not account for the B/Be variability. Temporal variations in slab fluid-flux alone are capable of reconciling all three chemical groups. 9. CONCLUSIONS We demonstrate here that since 7 Ma there have been significant changes in the relative volume of slab fluid-flux into the Tonga Arc in the vicinity of Ata, and that these variations have been much higher than any variability now seen along the strike of the Tonga-Kermadec Arc. Peak B/Be at 3 to 4 Ma may 3362 Peter D. Clift et al. represent either the effect of accelerated subduction due to regional plate readjustments at ⬃5 Ma (Cande et al., 1995) or the influence of Lau Basin rifting and the start of slab roll-back towards the east. A regional explanation of increased westward plate velocity would account for the peak in B/Be also seen in the Marianas at 3 to 4 Ma. The high peak B/Be values are not compatible with simple contamination of the mantle source by sediment, and require much of the boron to be derived by aqueous fluid from the AOC, an explanation compatible with the high positive ␦11B compositions. However, extreme isotopic fractionation within a cold, fast subduction zone is also required to explain the very high ␦11B values observed. Peaks in the volume of slab fluid-flux measured by B/Be do not reflect periods of preferential sediment subduction, but may indicate changes in the rate of plate convergence. The collision of Louisville Ridge with the Tonga Trench does not seem to affect the boron systematics. Latitudinal variation in the boron isotopic composition along the Tonga-Kermadec Arc can be linked to greater degrees of sediment involvement in petrogenesis close to New Zealand, but not to passage of the Louisville Ridge in northern Tonga. This pattern reflects the distinctive negative ␦11B composition of the continental-derived material, but the similarity of the positive boron isotopic character of the basement and volcaniclastic cover of the Louisville Ridge to with normal oceanic crust. Periods of increased slab flux (high B/Be) correlate with greater depletion in HFSEs and REEs, and suggest that flux of aqueous fluid from the AOC is the primary control on the degree of partial melting under the arc. Acknowledgments—We wish to thank Tony Ewart and David Tappin for their donation of modern volcanic specimens from the TongaKermadec Arc. PC thanks Bill Bryan, Jim Hawkins and Tony Ewart for first interesting him in the Tonga Arc. 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