Download Tracing the evolving flux from the subducting plate in the Tonga

Geochimica et Cosmochimica Acta, Vol. 65, No. 19, pp. 3347–3364, 2001
Copyright © 2001 Elsevier Science Ltd
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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).
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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
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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‰
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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
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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
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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. Support for EFR was provided
by French “Programme Lavoisier” Post-Doctoral Fellowship and J.
Steward Johnson Scholar from WHOI. We thank Achim Kopf and
Marc Chaussidon for their helpful reviews in improving this paper.
This is WHOI contribution 10502.
Associate editor: D. B. Dingwell
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