Abstract

New thermal ionization mass spectrometry U-series disequilibrium data are presented for 24 basaltic to dacitic glasses from active spreading centres in the back-arc Lau Basin (SW Pacific), together with additional inductively coupled plasma mass spectrometry trace element analyses and Sr–Nd–Pb isotope data. Valu Fa Ridge samples, adjacent to the arc front, have high U/Th and (230Th/238U) <1, implying a recent (≪350 ka) addition of a U-rich slab-derived fluid. The Valu Fa data can be combined with existing 230Th–238U data for the Central Tonga arc to infer a fluid addition event at ∼50 ka. The similar sources and time scales for fluid transfer beneath the Valu Fa Ridge and beneath the arc itself suggest that the Valu Fa Ridge is propagating into the arc-front region. Central Lau Basin samples, further behind the arc, have lower U/Th and (230Th/238U) ≥1, similar to typical mid-ocean ridge basalts (MORB). Within the Central Lau Basin, a water-rich subduction component is seen only in samples closest to the arc, and this fluid does not have the high-U/Th composition of the fluid at Valu Fa. Melt generation in the Central Lau Basin appears to be dominated by normal ridge-type processes, but the relatively low (230Th/238U) for these shallow ridges compared with global MORB could be a consequence of increased melt productivity as a result of the elevated water contents. The transition from 238U to 230Th excesses within the back-arc basin is not a smooth function of decreasing addition of a U-rich fluid moving away from the arc front, but also reflects the complex dynamics between two major mantle domains within the mantle wedge (‘Pacific’ beneath Valu Fa Ridge, ‘Indian’ beneath the Central Lau Basin).

INTRODUCTION

A perplexing aspect of back-arc magmatism is the exact role of the adjacent subduction zone in influencing the nature of melt generation. It is not clear to what extent the compositions of magmas erupted at back-arc spreading centres are dominated by processes common to typical mid-ocean ridges, and what additional complexities are brought by the supra-subduction zone setting. The wide compositional range observed in back-arc magmas, from mid-ocean ridge basalt (MORB)-like to arc-like, is generally attributed to progressive re-enrichment of a variably depleted mantle wedge source by slab-derived fluids (e.g. Saunders & Tarney, 1979; Pearce et al., 1984; Sinton & Fryer, 1987; Stolper & Newman, 1994; Pearce et al., 1995). Stolper & Newman (1994) further argued that the fluxing of the mantle source by H2O-rich fluids has an important control on the degree of melting in back-arc magmas. Important questions remain concerning the nature and transport of the slab-derived fluid, in particular about how quickly and how far it pervades the mantle wedge into the back-arc region, and how much it equilibrates with the mantle wedge during transport. It is also not clear how the processes of melt generation in back-arcs compare with those at other spreading centres such as mid-ocean ridges. U-series isotope data offer a unique opportunity to address these questions, as 238U–230Th isotope disequilibria record the effects of recent (<350 ka) U–Th fractionation and are sensitive to recent fluid inputs, mantle mineralogy, and the dynamics of mantle melting.

The processes that control 238U–230Th disequilibria in magmas from island-arc environments appear to be different from those in mid-ocean ridge settings (e.g. Newman et al., 1984; Gill & Williams, 1990; McDermott & Hawkesworth, 1991; Hawkesworth et al., 1997a). MORB magmas predominantly erupt with (230Th/238U) >1, and the excess 230Th is generally attributed to in-growth during partial melting initiated within the garnet stability field (e.g. Goldstein et al., 1991; Beattie, 1993; LaTourette et al., 1993; Lundstrom et al., 1995, 1998b; Bourdon et al., 1996b; Elliott, 1997). For island-arc magmas, 238U–230Th disequilibria are more common in rocks depleted in highly incompatible elements, and these samples mainly show (230Th/238U) <1. Such behaviour is consistent with recent addition of a U-rich slab-derived fluid to the mantle wedge source (e.g. Newman et al., 1984; Gill & Williams, 1990; McDermott & Hawkesworth, 1991; Condomines & Sigmarsson, 1993; Elliott et al., 1997; Hawkesworth et al., 1997a; Turner et al., 1997). It is not clear which of these two competing effects (i.e. recent fluid input vs source or melting processes) will dominate to control the sense of 238U–230Th disequilibrium in melts erupted at back-arc spreading centres. Presumably this will vary on both a local and a global scale, reflecting a variable influence from the subduction zone. There will be a complex interplay between several factors: the composition, amount, and timing of slab-fluid addition to the back-arc mantle, the fertility of the back-arc mantle, and the dynamics of mantle upwelling beneath the spreading centre.

The central Lau Basin is an ideal back-arc setting in which to investigate this problem because the active spreading centres are at varying distances from the arc front and trench (Fig. 1). Furthermore, geochemical studies have suggested that the added subduction component in the mantle source changes progressively with arc proximity (e.g. Pearce et al., 1995). In this paper, we present the first detailed study of U-series disequilibria in back-arc magmas, with new analyses on dredged glasses from the active spreading centres of the Lau Basin, to assess the timing and influence of variable addition of a volatile-rich subduction component on melting processes in the back-arc.

Fig. 1.

Tectonic map of the Tonga–Lau arc–back-arc system, after Hergt & Hawkesworth (1994) and Hawkins (1995a). ELSC, Eastern Lau Spreading Centre; ILSC, Intermediate Lau Spreading Centre; CLSC, Central Lau Spreading Centre; NLSC, Northern Lau Spreading Centre; PR, Peggy Ridge; MTJ, Mangatolu triple junction; NF, Niuafo’ou; N, Niuatoputapu; TA, Tafahi; FO, Fonualei; LA, Late; ME, Metis Shoal; KA, Kao; TO, Tofua; F, Falcon; HH, Hunga Ha’apai. ▵, recent submarine arc-front volcanoes. Grey shading indicates areas with water depths <2 km. The Kermadec arc is the lateral continuation of the Tonga arc, south of 23°S. •, location of drill sites from the ODP Leg 135. Horizontally striped shaded regions represent new crust formed by true sea-floor spreading on the ELSC and CLSC. The two boxes indicate the locations of the studied samples.

Fig. 1.

Tectonic map of the Tonga–Lau arc–back-arc system, after Hergt & Hawkesworth (1994) and Hawkins (1995a). ELSC, Eastern Lau Spreading Centre; ILSC, Intermediate Lau Spreading Centre; CLSC, Central Lau Spreading Centre; NLSC, Northern Lau Spreading Centre; PR, Peggy Ridge; MTJ, Mangatolu triple junction; NF, Niuafo’ou; N, Niuatoputapu; TA, Tafahi; FO, Fonualei; LA, Late; ME, Metis Shoal; KA, Kao; TO, Tofua; F, Falcon; HH, Hunga Ha’apai. ▵, recent submarine arc-front volcanoes. Grey shading indicates areas with water depths <2 km. The Kermadec arc is the lateral continuation of the Tonga arc, south of 23°S. •, location of drill sites from the ODP Leg 135. Horizontally striped shaded regions represent new crust formed by true sea-floor spreading on the ELSC and CLSC. The two boxes indicate the locations of the studied samples.

THE LAU BASIN REGION AND SAMPLE DETAILS

The Lau Basin is a triangular-shaped, actively spreading back-arc basin, situated behind the Tonga arc in the SW Pacific (Fig. 1). Two comprehensive up-to-date reviews of the geology of the Lau Basin have been published recently that incorporate new results from extensive geophysical surveys, dredged material and Ocean Drilling Program (ODP) drilling (Hawkins, 1995a, 1995b), and the following summary is largely based on these reviews. Samples selected for this study come from the central and southern parts of the Lau Basin, which have a relatively simple tectonic arrangement. Here, the basin initially opened solely by crustal extension and rifting. The first new oceanic crust formed at ∼5·5–5·0 Ma as sea-floor spreading began on a southward-propagating rift, the Eastern Lau Spreading Centre (ELSC). A younger rift, the Central Lau Spreading Centre (CLSC), was also initiated at the Peggy Ridge transform fault but further from the arc at ∼1·5–1·2 Ma, and it is propagating southwards at the expense of the ELSC. The two rifts overlap at ∼19·3°S and there is a small ‘relay’ spreading segment, the Intermediate Lau Spreading Centre (ILSC), between them. Thus, the active spreading centres are offset and become closer to the active Tonga arc-front volcanoes towards the south: the southern end of the ELSC, also known as the Valu Fa Ridge (VFR), is only ∼40 km from Ata volcano. Sea-floor spreading rates within the Lau Basin increase towards the north, consistent with the inverted triangular shape of the basin, and estimates derived from interpretation of magnetic lineations are 65 mm/yr at 21°S and 90 mm/yr at 18°S (full spreading rate: Taylor et al., 1996). Spreading rate estimates based on global positioning satellite (GPS) geodetic measurements (Bevis et al., 1995) are about twice these values, but this apparent discrepancy has recently been resolved in favour of the lower values by the recognition of a Niuafo’ou microplate between the Australia and Tonga plates north of 19°20′S (Zellmer & Taylor, 1999).

The northern part of the basin (north of 17°N) was not included in the present study for two reasons: (1) it has a complicated tectonic configuration (Fig. 1); (2) many of the lavas in this region, including the nearby arc-front islands (Tafahi and Niuatoputapu), show evidence for the presence of additional compositional components derived from the Samoan plume to the NE and also mobilized from volcaniclastic sediments on the subducting slab associated with the Louisville seamounts (Volpe et al., 1988; Hawkins, 1995a; Regelous et al., 1997; Turner & Hawkesworth, 1997, 1998; Turner et al., 1997; Wendt et al., 1997; Ewart et al., 1998). These enriched components are not evident in magmas erupted further south (i.e. the Central Lau Basin, Valu Fa Ridge and the main Tonga arc), or in the northern rear-arc island of Niuafo’ou.

A distinctive feature of the composition of the lavas from the Lau spreading centres is the presence, in addition to basalts, of highly fractionated samples (ferrobasalts, andesites: e.g. Vallier et al., 1991; Pearce et al., 1995), which we have also analysed. These evolved rocks are found near the propagating rift tips of both the ELSC–VFR and the CLSC, and similar rocks are often found on propagating segments of the global mid-ocean ridge system (Christie & Sinton, 1981).

The selection of samples for U–Th–Ra analysis was based both on the availability of suitable quantities of fresh glass, hopefully of ‘zero-age’ (<10 ka), dredged from the active spreading axes, and on obtaining a reasonable geographical coverage from the different spreading centres. Samples from the central part of the Lau Basin (CLSC, ILSC and ELSC) came from the R.S.S. Charles Darwin (CD33) cruise that dredged the neovolcanic zone (identified from GLORIA images) between 18°50′ and 20°35′S (Pearce et al., 1995), with a few additional samples from the northern CLSC (18°34′–18°44′S) from the R.V. Sonne 1987 (SO48) cruise (Sunkel, 1990). Samples from the Valu Fa Ridge came from dredged material from the R.V. Lee 1984 cruise near 22°20′S (Jenner et al., 1987; Vallier et al., 1991) and from the R.V. Sonne 1984 (SO35) and 1987 (SO48) cruises (Boespflug et al., 1990; Sunkel, 1990; Bach & Niedermann, 1998; Bach et al., 1998). The locations of the studied samples (latitude, longitude, depth) are given in Table 1. Additional trace element and radiogenic (Sr–Nd–Pb) isotope analyses were carried out (Tables 1 and 2) to complete the existing compositional data available on these samples (Jenner et al., 1987; Loock et al., 1990; Sunkel, 1990; Vallier et al., 1991; Macpherson & Mattey, 1994, 1998; Pearce et al., 1995; Bach & Niedermann, 1998; Bach et al., 1998).

Table 1:

Major and trace element data on glasses from active spreading centres in the back-arc Lau Basin, SW Pacific

Sample: 92KD1 90KD1 84KD1 114KD 133GA 61GC 55GC L-1-11 L-2-5 22-6-1 20-5-2 
Cruise: SO35 SO35 SO35 SO48 SO48 SO48 SO48 Lee’84 Lee’84 CD33 CD33 
Lat. (°S): 22·374 22·325 22·140 22·186 22·178 21·822 21·479 22·315 22·331 19·688 19·488 
Long. (°W): 176·673 176·656 176·610 176·555 176·617 176·442 176·380 176·650 176·656 175·993 175·963 
Av. depth            
(m): 1975 1880 1723 1505 1620 1760 1793 2014 1903 2730 2640 
Location: VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
ELSC
 
ELSC
 
SiO2  53·41  54·88  50·36  50·91  56·70  51·01  53·30  56·28  54·77  51·40  51·23 
TiO2   1·55   1·61   0·76   0·58   1·44   0·68   1·43   1·46   1·74   1·19   0·96 
Al2O3  14·60  14·49  15·92  14·80  14·76  15·36  14·96  14·21  14·18  14·53  14·89 
Fe2O3(t)  12·93  12·37   9·85  11·11  11·01   9·26  11·95  12·61  13·53  11·84  10·93 
MnO   0·19   0·19   0·16   0·18   0·19   0·16   0·18   0·21   0·21   0·19   0·18 
MgO   3·84   3·27   7·31   7·18   2·96   7·79   4·39   3·23   3·57   7·06   7·93 
CaO   8·30   7·60  12·61  12·36   7·34  12·63   8·75   7·39   7·81  11·08  11·77 
Na2  3·08   3·44   1·55   1·30   3·57   1·51   3·13   3·27   3·08   2·39   1·99 
K2  0·35   0·41   0·16   0·28   0·44   0·23   0·35   0·49   0·41   0·06   0·05 
P2O5   0·18   0·21   0·06   0·07   0·25   0·08   0·18   0·20   0·20   0·09   0·06 
H2O+   1·58   1·53   1·26   1·23   1·36   1·30   1·38   1·30   1·22  −0·47  −0·18 
Sc  33·4  30·8  39·6  43·1  23·2  37·8  29·5  30·8  35·0  43·9  41·7 
379 259 311 280 149 260 312 421 483 388 373 
Cr   2   2  99  40   6 181  14   3   3  81 153 
Co  34·0  28·2  36·5  41·8  22·0  39·0  31·8  30·2  36·9  46·4  46·0 
Ni  12   2  54  38   6  94  15   7   9  54  67 
Cu  59  37  93 112  33 103  46  69  67  97  82 
Zn 101 102  65  61 126  59  90  98 105  90  83 
Ga  17·7  17·6  13·8  12·8  17·2  13·2  17·2  18·7  19·0  16·7  15·7 
Rb   6·43   7·25   2·59   3·74   7·56   4·44   6·97   7·67   6·68   1·58   0·92 
Sr 160 164 126 160 164 123 149 162 165  83  59 
 33·4  37·6  16·8  13·6  38·1  17·4  34·7  36·7  33·9  32·7  28·3 
Zr  64·8  77·3  27·5  19·6  78·5  28·6  77·6  78·9  67·7  65·4  49·6 
Nb   0·91   1·14   0·37   0·23   1·10   0·47   1·17   1·11   0·95   1·72   0·88 
Cs   0·16   0·19   0·09   0·14   0·20   0·10   0·13   0·18   0·16   0·01   0·01 
Ba  92·5 101·1  48·3  81·7 109·6  55·7  75·6 101  96  13·4   7·3 
La   3·93   4·68   1·51   1·70   4·84   2·07   4·13   4·38   3·91   2·15   1·42 
Ce  10·76  12·73   4·21   4·28  13·42   5·46  11·71  12·2  10·9   6·82   4·82 
Pr   1·77   2·08   0·72   0·70   2·15   0·87   1·92   2·07   1·86   1·25   0·94 
Nd   9·40  11·00   3·82   3·97  11·60   4·55  10·01  11·3  10·2   7·33   5·56 
Sm   3·27   3·84   1·46   1·29   4·00   1·58   3·46   3·69   3·33   2·62   2·16 
Eu   1·22   1·40   0·59   0·53   1·44   0·63   1·25   1·31   1·25   1·00   0·80 
Gd   4·46   5·16   2·14   1·89   5·27   2·31   4·82   4·91   4·52   4·03   3·28 
Tb   0·80   0·61   0·39   0·32   0·93   0·41   0·85   0·88   0·83   0·75   0·63 
Dy   5·25   5·96   2·59   2·16   6·10   2·74   5·52   5·84   5·42   5·12   4·38 
Ho   1·13   1·31   0·57   0·48   1·34   0·60   1·20   1·28   1·17   1·12   0·98 
Er   3·30   3·81   1·64   1·38   3·81   1·76   3·52   3·75   3·39   3·27   2·89 
Tm   0·52   0·60   0·26   0·21   0·61   0·28   0·55   0·59   0·53   0·52   0·46 
Yb   3·30   3·80   1·67   1·39   3·87   1·74   3·48   3·77   3·45   3·29   2·97 
Lu   0·51   0·60   0·26   0·22   0·59   0·28   0·54   0·60   0·55   0·52   0·48 
Hf   1·95   2·32   0·85   0·66   2·40   0·89   2·24   2·38   2·07   1·95   1·52 
Ta   0·075   0·091   0·040   0·025   0·078   0·046   0·090   0·082   0·075   0·123   0·069 
Pb   1·19   1·22   0·66   0·86   1·23   0·58   0·94   1·22   1·14   0·443   0·283 
Th   0·355   0·432   0·133   0·199   0·439   0·196   0·335   0·417   0·363   0·137   0·072 
  0·155   0·178   0·059   0·097   0·192   0·081   0·134   0·177   0·164   0·041   0·024 
Sample: 92KD1 90KD1 84KD1 114KD 133GA 61GC 55GC L-1-11 L-2-5 22-6-1 20-5-2 
Cruise: SO35 SO35 SO35 SO48 SO48 SO48 SO48 Lee’84 Lee’84 CD33 CD33 
Lat. (°S): 22·374 22·325 22·140 22·186 22·178 21·822 21·479 22·315 22·331 19·688 19·488 
Long. (°W): 176·673 176·656 176·610 176·555 176·617 176·442 176·380 176·650 176·656 175·993 175·963 
Av. depth            
(m): 1975 1880 1723 1505 1620 1760 1793 2014 1903 2730 2640 
Location: VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
VFR
 
ELSC
 
ELSC
 
SiO2  53·41  54·88  50·36  50·91  56·70  51·01  53·30  56·28  54·77  51·40  51·23 
TiO2   1·55   1·61   0·76   0·58   1·44   0·68   1·43   1·46   1·74   1·19   0·96 
Al2O3  14·60  14·49  15·92  14·80  14·76  15·36  14·96  14·21  14·18  14·53  14·89 
Fe2O3(t)  12·93  12·37   9·85  11·11  11·01   9·26  11·95  12·61  13·53  11·84  10·93 
MnO   0·19   0·19   0·16   0·18   0·19   0·16   0·18   0·21   0·21   0·19   0·18 
MgO   3·84   3·27   7·31   7·18   2·96   7·79   4·39   3·23   3·57   7·06   7·93 
CaO   8·30   7·60  12·61  12·36   7·34  12·63   8·75   7·39   7·81  11·08  11·77 
Na2  3·08   3·44   1·55   1·30   3·57   1·51   3·13   3·27   3·08   2·39   1·99 
K2  0·35   0·41   0·16   0·28   0·44   0·23   0·35   0·49   0·41   0·06   0·05 
P2O5   0·18   0·21   0·06   0·07   0·25   0·08   0·18   0·20   0·20   0·09   0·06 
H2O+   1·58   1·53   1·26   1·23   1·36   1·30   1·38   1·30   1·22  −0·47  −0·18 
Sc  33·4  30·8  39·6  43·1  23·2  37·8  29·5  30·8  35·0  43·9  41·7 
379 259 311 280 149 260 312 421 483 388 373 
Cr   2   2  99  40   6 181  14   3   3  81 153 
Co  34·0  28·2  36·5  41·8  22·0  39·0  31·8  30·2  36·9  46·4  46·0 
Ni  12   2  54  38   6  94  15   7   9  54  67 
Cu  59  37  93 112  33 103  46  69  67  97  82 
Zn 101 102  65  61 126  59  90  98 105  90  83 
Ga  17·7  17·6  13·8  12·8  17·2  13·2  17·2  18·7  19·0  16·7  15·7 
Rb   6·43   7·25   2·59   3·74   7·56   4·44   6·97   7·67   6·68   1·58   0·92 
Sr 160 164 126 160 164 123 149 162 165  83  59 
 33·4  37·6  16·8  13·6  38·1  17·4  34·7  36·7  33·9  32·7  28·3 
Zr  64·8  77·3  27·5  19·6  78·5  28·6  77·6  78·9  67·7  65·4  49·6 
Nb   0·91   1·14   0·37   0·23   1·10   0·47   1·17   1·11   0·95   1·72   0·88 
Cs   0·16   0·19   0·09   0·14   0·20   0·10   0·13   0·18   0·16   0·01   0·01 
Ba  92·5 101·1  48·3  81·7 109·6  55·7  75·6 101  96  13·4   7·3 
La   3·93   4·68   1·51   1·70   4·84   2·07   4·13   4·38   3·91   2·15   1·42 
Ce  10·76  12·73   4·21   4·28  13·42   5·46  11·71  12·2  10·9   6·82   4·82 
Pr   1·77   2·08   0·72   0·70   2·15   0·87   1·92   2·07   1·86   1·25   0·94 
Nd   9·40  11·00   3·82   3·97  11·60   4·55  10·01  11·3  10·2   7·33   5·56 
Sm   3·27   3·84   1·46   1·29   4·00   1·58   3·46   3·69   3·33   2·62   2·16 
Eu   1·22   1·40   0·59   0·53   1·44   0·63   1·25   1·31   1·25   1·00   0·80 
Gd   4·46   5·16   2·14   1·89   5·27   2·31   4·82   4·91   4·52   4·03   3·28 
Tb   0·80   0·61   0·39   0·32   0·93   0·41   0·85   0·88   0·83   0·75   0·63 
Dy   5·25   5·96   2·59   2·16   6·10   2·74   5·52   5·84   5·42   5·12   4·38 
Ho   1·13   1·31   0·57   0·48   1·34   0·60   1·20   1·28   1·17   1·12   0·98 
Er   3·30   3·81   1·64   1·38   3·81   1·76   3·52   3·75   3·39   3·27   2·89 
Tm   0·52   0·60   0·26   0·21   0·61   0·28   0·55   0·59   0·53   0·52   0·46 
Yb   3·30   3·80   1·67   1·39   3·87   1·74   3·48   3·77   3·45   3·29   2·97 
Lu   0·51   0·60   0·26   0·22   0·59   0·28   0·54   0·60   0·55   0·52   0·48 
Hf   1·95   2·32   0·85   0·66   2·40   0·89   2·24   2·38   2·07   1·95   1·52 
Ta   0·075   0·091   0·040   0·025   0·078   0·046   0·090   0·082   0·075   0·123   0·069 
Pb   1·19   1·22   0·66   0·86   1·23   0·58   0·94   1·22   1·14   0·443   0·283 
Th   0·355   0·432   0·133   0·199   0·439   0·196   0·335   0·417   0·363   0·137   0·072 
  0·155   0·178   0·059   0·097   0·192   0·081   0·134   0·177   0·164   0·041   0·024 
Sample: 41-2-1 46GC 18GA2 42GC 10-1-3 11-2-1 13-2 12-5-2 15-1-1 X-108 QL 
Cruise: CD33 SO48 SO48 SO48 CD33 CD33 CD33 CD33 CD33 internal std. 10σ 
Lat (°S): 19·277 18·735 18·595 18·576 19·142 19·115 18·973 18·922 18·883 boninite (ppm) 
Long (°W): 176·173 176·511 176·445 176·428 176·528 176·553 176·567 176·553 176·562 Chichijima
 
 
Av. depth          Mean 2 SD  
(m): 3050 2280 2260 2253 2260 2250 2255 2345 2295 (ppm) (%)  
Location: ILSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 

 

 

 
SiO2  50·50  50·03  50·23  49·72  51·50  51·00  60·38  58·75  50·53 — — — 
TiO2  0·82   1·37   1·46   1·08   2·53   2·50   1·25   1·56   1·15 — — — 
Al2O3  16·18  13·56  13·84  14·37  12·27  12·08  12·73  12·68  15·88 — — — 
Fe2O3(t)   8·86  13·14  12·53  11·15  18·39  18·91  12·33  12·55  10·80 — — — 
MnO   0·15   0·20   0·19   0·18   0·27   0·28   0·21   0·20   0·17 — — — 
MgO   8·11   7·05   7·10   8·19   3·30   4·70   1·42   2·63   7·30 — — — 
CaO  12·82  11·65  11·30  12·62   7·92   8·66   5·37   6·45  12·21 — — — 
Na2  1·84   2·56   2·74   2·28   3·34   2·99   4·31   4·21   2·43 — — — 
K2  0·06   0·06   0·06   0·04   0·17   0·09   0·39   0·31   0·04 — — — 
P2O5   0·06   0·10   0·11   0·07   0·43   0·20   0·38   0·33   0·07 — — — 
H2O+  −0·23   0·29   0·44   0·30  −0·47  −0·82   1·18   1·18  −0·53 — — — 
Sc  40·3  40·5  42·1  44·0  35·7  46·5  18·9  29·1  44·6  33·6   6·4   7·2 
286 363 351 313 291 652  57 228 338 211   5·8   2·4 
Cr 426  73 143 305  10   8   6   8 325 455   9·8 11·6 
Co  43·6  46·5  44·7  46·2  39·9  57·3  17·8  31·1  41·8  37   5·0   0·07 
Ni 104  65  66 106  25  31  19  13  73 127   2·6   0·50 
Cu  89 105  95 116  55  97  37  49  99  82   2·7   1·37 
Zn  62 102  95  81 178 166 160 136  74  61   3·4   6·9 
Ga  14·4  16·6  16·8  15·7  34·6  23·4  27·4  23·7  16·1   9·1   4·2   0·06 
Rb   1·38   0·73   0·74   0·65   3·85   1·84   8·67   5·52   0·72  12·9   2·2   0·03 
Sr  88  75  85  80  86  84  89  86  94  90   3·0   0·07 
 20·6  36·0  39·4  30·2 116  71·9 171 132  29·8   4·3   3·4   0·008 
Zr  38·9  69·6  82·4  54·6 301 159 584 418  61·8  24·5   6·4   0·06 
Nb   1·33   1·27   1·28   0·95   5·84   3·09   8·76   6·39   1·12   0·52   9·8   0·016 
Cs   0·01   0·01   0·01   0·01   0·04   0·02   0·10   0·06   0·01   0·71   3·0   0·00 
Ba  15·2   9·4   8·2   6·9  30·4  17·2  62·0  38·9   7·7  36   5·4   0·22 
La   1·60   2·06   2·39   1·58   8·85   4·59  16·8  12·6   1·73   0·91   2·8   0·004 
Ce   4·69   7·02   8·18   5·40  29·3  15·4  52·9  40·0   5·91   1·74   3·0   0·006 
Pr   0·83   1·29   1·50   1·01   5·37   2·91   9·28   7·07   1·13   0·28   4·2   0·003 
Nd   4·84   7·44   8·63   5·87  30·8  17·0  50·7  38·8   6·82   1·24   5·0   0·013 
Sm   1·72   3·02   3·42   2·42  10·79   6·22  16·45  12·69   2·54   0·35   5·0   0·013 
Eu   0·68   1·13   1·25   0·94   3·13   2·12   4·09   3·13   0·98   0·12   6·0   0·004 
Gd   2·50   4·53   5·01   3·74  14·90   9·04  22·08  17·06   3·68   0·43   5·5   0·014 
Tb   0·47   0·84   0·93   0·70   2·76   1·70   4·07   3·14   0·71   0·083   9·6   0·003 
Dy   3·21   5·70   6·24   4·69  18·36  11·44  26·98  20·81   4·78   0·57   3·8   0·008 
Ho   0·70   1·25   1·37   1·02   4·00   2·49   5·90   4·56   1·05   0·14   3·4   0·002 
Er   2·05   3·67   3·95   2·98  11·63   7·27  17·38  13·33   3·04   0·46   3·8   0·004 
Tm   0·32   0·57   0·61   0·46   1·82   1·12   2·77   2·10   0·47   0·085   7·0   0·002 
Yb   2·08   3·61   3·89   2·94  11·58   7·24  17·79  13·53   3·01   0·62   2·1   0·004 
Lu   0·33   0·56   0·59   0·45   1·81   1·14   2·78   2·13   0·47   0·114   9·2   0·003 
Hf   1·16   2·06   2·41   1·62   8·56   4·61  16·46  11·91   1·80   0·66   9·4   0·011 
Ta   0·094   0·104   0·105   0·080   0·417   0·221   0·608   0·457   0·089   0·043  20   0·009 
Pb   0·274   0·32   0·49   0·30   1·11   0·606   2·54   1·32   0·348   1·76   9·8   0·12 
Th   0·113   0·108   0·113   0·075   0·509   0·244   1·20   0·966   0·081   0·123   5·4   0·002 
  0·035   0·037   0·042   0·024   0·166   0·085   0·406   0·350   0·028   0·122   5·5   0·003 
Sample: 41-2-1 46GC 18GA2 42GC 10-1-3 11-2-1 13-2 12-5-2 15-1-1 X-108 QL 
Cruise: CD33 SO48 SO48 SO48 CD33 CD33 CD33 CD33 CD33 internal std. 10σ 
Lat (°S): 19·277 18·735 18·595 18·576 19·142 19·115 18·973 18·922 18·883 boninite (ppm) 
Long (°W): 176·173 176·511 176·445 176·428 176·528 176·553 176·567 176·553 176·562 Chichijima
 
 
Av. depth          Mean 2 SD  
(m): 3050 2280 2260 2253 2260 2250 2255 2345 2295 (ppm) (%)  
Location: ILSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 
CLSC
 

 

 

 
SiO2  50·50  50·03  50·23  49·72  51·50  51·00  60·38  58·75  50·53 — — — 
TiO2  0·82   1·37   1·46   1·08   2·53   2·50   1·25   1·56   1·15 — — — 
Al2O3  16·18  13·56  13·84  14·37  12·27  12·08  12·73  12·68  15·88 — — — 
Fe2O3(t)   8·86  13·14  12·53  11·15  18·39  18·91  12·33  12·55  10·80 — — — 
MnO   0·15   0·20   0·19   0·18   0·27   0·28   0·21   0·20   0·17 — — — 
MgO   8·11   7·05   7·10   8·19   3·30   4·70   1·42   2·63   7·30 — — — 
CaO  12·82  11·65  11·30  12·62   7·92   8·66   5·37   6·45  12·21 — — — 
Na2  1·84   2·56   2·74   2·28   3·34   2·99   4·31   4·21   2·43 — — — 
K2  0·06   0·06   0·06   0·04   0·17   0·09   0·39   0·31   0·04 — — — 
P2O5   0·06   0·10   0·11   0·07   0·43   0·20   0·38   0·33   0·07 — — — 
H2O+  −0·23   0·29   0·44   0·30  −0·47  −0·82   1·18   1·18  −0·53 — — — 
Sc  40·3  40·5  42·1  44·0  35·7  46·5  18·9  29·1  44·6  33·6   6·4   7·2 
286 363 351 313 291 652  57 228 338 211   5·8   2·4 
Cr 426  73 143 305  10   8   6   8 325 455   9·8 11·6 
Co  43·6  46·5  44·7  46·2  39·9  57·3  17·8  31·1  41·8  37   5·0   0·07 
Ni 104  65  66 106  25  31  19  13  73 127   2·6   0·50 
Cu  89 105  95 116  55  97  37  49  99  82   2·7   1·37 
Zn  62 102  95  81 178 166 160 136  74  61   3·4   6·9 
Ga  14·4  16·6  16·8  15·7  34·6  23·4  27·4  23·7  16·1   9·1   4·2   0·06 
Rb   1·38   0·73   0·74   0·65   3·85   1·84   8·67   5·52   0·72  12·9   2·2   0·03 
Sr  88  75  85  80  86  84  89  86  94  90   3·0   0·07 
 20·6  36·0  39·4  30·2 116  71·9 171 132  29·8   4·3   3·4   0·008 
Zr  38·9  69·6  82·4  54·6 301 159 584 418  61·8  24·5   6·4   0·06 
Nb   1·33   1·27   1·28   0·95   5·84   3·09   8·76   6·39   1·12   0·52   9·8   0·016 
Cs   0·01   0·01   0·01   0·01   0·04   0·02   0·10   0·06   0·01   0·71   3·0   0·00 
Ba  15·2   9·4   8·2   6·9  30·4  17·2  62·0  38·9   7·7  36   5·4   0·22 
La   1·60   2·06   2·39   1·58   8·85   4·59  16·8  12·6   1·73   0·91   2·8   0·004 
Ce   4·69   7·02   8·18   5·40  29·3  15·4  52·9  40·0   5·91   1·74   3·0   0·006 
Pr   0·83   1·29   1·50   1·01   5·37   2·91   9·28   7·07   1·13   0·28   4·2   0·003 
Nd   4·84   7·44   8·63   5·87  30·8  17·0  50·7  38·8   6·82   1·24   5·0   0·013 
Sm   1·72   3·02   3·42   2·42  10·79   6·22  16·45  12·69   2·54   0·35   5·0   0·013 
Eu   0·68   1·13   1·25   0·94   3·13   2·12   4·09   3·13   0·98   0·12   6·0   0·004 
Gd   2·50   4·53   5·01   3·74  14·90   9·04  22·08  17·06   3·68   0·43   5·5   0·014 
Tb   0·47   0·84   0·93   0·70   2·76   1·70   4·07   3·14   0·71   0·083   9·6   0·003 
Dy   3·21   5·70   6·24   4·69  18·36  11·44  26·98  20·81   4·78   0·57   3·8   0·008 
Ho   0·70   1·25   1·37   1·02   4·00   2·49   5·90   4·56   1·05   0·14   3·4   0·002 
Er   2·05   3·67   3·95   2·98  11·63   7·27  17·38  13·33   3·04   0·46   3·8   0·004 
Tm   0·32   0·57   0·61   0·46   1·82   1·12   2·77   2·10   0·47   0·085   7·0   0·002 
Yb   2·08   3·61   3·89   2·94  11·58   7·24  17·79  13·53   3·01   0·62   2·1   0·004 
Lu   0·33   0·56   0·59   0·45   1·81   1·14   2·78   2·13   0·47   0·114   9·2   0·003 
Hf   1·16   2·06   2·41   1·62   8·56   4·61  16·46  11·91   1·80   0·66   9·4   0·011 
Ta   0·094   0·104   0·105   0·080   0·417   0·221   0·608   0·457   0·089   0·043  20   0·009 
Pb   0·274   0·32   0·49   0·30   1·11   0·606   2·54   1·32   0·348   1·76   9·8   0·12 
Th   0·113   0·108   0·113   0·075   0·509   0·244   1·20   0·966   0·081   0·123   5·4   0·002 
  0·035   0·037   0·042   0·024   0·166   0·085   0·406   0·350   0·028   0·122   5·5   0·003 

VFR, Valu Fa Ridge; ELSC, Eastern Lau Spreading Centre; ILSC, Intermediate Lau Spreading Centre; CLSC, Central Lau Spreading Centre. Major element data from Sunkel (1990), Vallier et al. (1991), Pearce et al. (1995) and Bach et al. (1998). New trace element analyses by ICP-MS at the University of Durham, UK, using a Perkin Elmer Sciex Elan 6000. X-108 is a depleted boninite sample from Chichijima, used as an internal laboratory standard. The mean and 2σ between-run RSD in per cent are given (n = 28), plus an estimate of the quantification limit (QL) of the ICP-MS technique at Durham (10σ background). A compilation of data for X-108 analysed previously by ICP-MS at Durham on the old VG PlasmaQuad1 instrument has been given by Peate & Pearce (1998).

Table 2:

Radiogenic isotope data for back-arc glasses from the Lau Basin, SW Pacific

Sample Location
 
87Sr/86Sr
 
143Nd/144Nd
 
206Pb/204Pb
 
207Pb/204Pb
 
208Pb/204Pb
 
92KD1* VFR 0·703312 0·513053 18·668 15·558 38·367 
90KD1* VFR 0·703287 0·513046 18·658 15·554 38·355 
84KD1* VFR 0·703274 0·513049 18·656 15·553 38·358 
114KD* VFR 0·703389 0·513035 18·681 15·543 38·313 
133GA* VFR 0·703293 0·513053 18·572 15·557 38·370 
61GC* VFR 0·703306 0·513036 18·712 15·547 38·373 
55GC* VFR 0·703194 0·513052 18·622 15·556 38·356 
L-1 (av.)* VFR 0·703308 0·513034 18·639 15·550 38·312 
L-2 (av.)* VFR 0·703320 0·513036 18·658 15·554 38·352 
22-6-1 ELSC 0·703188 0·513065 18·216 15·494 38·054 
20-5-2 ELSC 0·703254 0·513084 18·223 15·511 38·092 
41-2-1 ILSC 0·703383 0·513003 18·263 15·463 38·015 
10-1-3 CLSC 0·703242 0·513137 18·088 15·462 37·859 
11-2-1 CLSC 0·703194 0·513121 18·119 15·472 37·911 
13-2 CLSC 0·703280 0·513010 18·149 15·475 37·951 
12-5-2 CLSC 0·703299 0·513110 18·156 15·474 37·948 
15-1-1 CLSC 0·703186 0·513107 18·161 15·495 37·973 
46GC CLSC — 0·513130 — — — 
18GA2* CLSC 0·703094 0·513102 18·099 15·476 37·928 
42GC* CLSC 0·703176 0·513090 18·151 15·475 37·966 
Sample Location
 
87Sr/86Sr
 
143Nd/144Nd
 
206Pb/204Pb
 
207Pb/204Pb
 
208Pb/204Pb
 
92KD1* VFR 0·703312 0·513053 18·668 15·558 38·367 
90KD1* VFR 0·703287 0·513046 18·658 15·554 38·355 
84KD1* VFR 0·703274 0·513049 18·656 15·553 38·358 
114KD* VFR 0·703389 0·513035 18·681 15·543 38·313 
133GA* VFR 0·703293 0·513053 18·572 15·557 38·370 
61GC* VFR 0·703306 0·513036 18·712 15·547 38·373 
55GC* VFR 0·703194 0·513052 18·622 15·556 38·356 
L-1 (av.)* VFR 0·703308 0·513034 18·639 15·550 38·312 
L-2 (av.)* VFR 0·703320 0·513036 18·658 15·554 38·352 
22-6-1 ELSC 0·703188 0·513065 18·216 15·494 38·054 
20-5-2 ELSC 0·703254 0·513084 18·223 15·511 38·092 
41-2-1 ILSC 0·703383 0·513003 18·263 15·463 38·015 
10-1-3 CLSC 0·703242 0·513137 18·088 15·462 37·859 
11-2-1 CLSC 0·703194 0·513121 18·119 15·472 37·911 
13-2 CLSC 0·703280 0·513010 18·149 15·475 37·951 
12-5-2 CLSC 0·703299 0·513110 18·156 15·474 37·948 
15-1-1 CLSC 0·703186 0·513107 18·161 15·495 37·973 
46GC CLSC — 0·513130 — — — 
18GA2* CLSC 0·703094 0·513102 18·099 15·476 37·928 
42GC* CLSC 0·703176 0·513090 18·151 15·475 37·966 

*Previously published analyses (Loock et al., 1990; Vallier et al., 1991; Bach et al., 1998). New Sr–Nd–Pb isotope analyses were carried out at the Open University, UK, using the methods outlined by Hergt & Hawkesworth (1994); as these were carried out at the same time, details of standards and analytical precision given by Hergt & Hawkesworth (1994) are applicable to the new data presented here.

† Analyses on compositionally identical glasses from same dredge haul (see Pearce et al., 1995): 11-2-1 (actually 11-1-1), 13-2 (actually 13-1-1), 12-5-2 (actually 12-5-3), 15-1-1 (actually 15-1-4).

The samples were glass chips, hand-picked under a binocular microscope to avoid fragments with any visible Mn-oxide or alteration coatings, that were then leached using a H2O2–HCl mixture before digestion (Bourdon et al., 1996a). The procedures for sample digestion and chemical extraction of U, Th and Ra as followed at the Open University have been given by Turner et al. (1997, 2000). All isotopic measurements were performed on a Finnigan MAT262 mass spectrometer equipped with a retarding potential quadrupole (RPQ-II) to provide the high abundance sensitivity capability necessary for high-precision measurements of 230Th. This system and the mass spectrometric procedures for U and Th analysis have been described by van Calsteren & Schwieters (1995). The external reproducibility on 230Th/232Th is ∼1·0% (2 SD), based on repeated measurements of an ‘in-house’ Th standard solution (Th‘U’std) that has 230Th/232Th of about 6 × 10−6 (van Calsteren & Schwieters, 1995), similar to that of the measured Lau Basin glasses. The procedure used for measuring Ra is similar to that outlined by Lundstrom et al. (1998a), with particular care taken to monitor organic interferences (Turner et al., 2000).

RESULTS

U, Th and Ra abundances and isotope ratios for 24 Lau Basin samples are reported in Table 3. All samples were dredged from the inferred axial spreading ridges, except for 114KD and 61GC, which are from near-axis seamounts on the Valu Fa Ridge (Sunkel, 1990). The relatively high spreading rates in the Lau Basin (65–90 mm/yr: Taylor et al., 1996) make it likely that the on-axis samples are younger than 10 ka such that no correction is required for post-eruptive decay of 230Th. This was checked for most samples by using the 226Ra–230Th decay scheme. The short half-life of 226Ra (1600 yr) means that any sample with 226Ra–230Th disequilibrium must be <8 ka. With a few exceptions, the samples have (234U/238U) values within 1% of secular equilibrium, indicating no post-eruption seawater alteration. The U–Th isotope results are presented on a (230Th/232Th) vs (238U/232Th) equiline diagram in Fig. 2. There is a clear contrast between the samples from the Valu Fa Ridge and those from the Central Lau Basin (ELSC, ILSC and CLSC) to the north, and so they will be described separately. The Valu Fa Ridge samples all have relatively high (238U/232Th) (1·28–1·37) and ‘arc-like’ (230Th/238U) <1, whereas the Central Lau Basin samples all have lower (238U/232Th) (0·98–1·15) and ‘MORB-like’ (230Th/238U) ≥1.

Table 3:

U-series disequilibria data for back-arc glasses from the Lau Basin, SW Pacific

Sample Location Th (238U/232Th) (234U/238U) (230Th/232Th) (230Th/238U) 226Ra (226Ra/230Th) 
 
 
(ppb)
 
(ppb)
 

 

 

 

 
(fg/g)
 

 
92KD1 VFR  349 156 1·355±06 1·011±08 1·213±12 0·895±10 — — 
90KD1 VFR  401 178 1·348±08 0·993±10 1·175±10 0·872±08 — — 
84KD1 VFR  138  61·3 1·347±04 1·019±11 1·170±08 0·869±07 — — 
114KD VFR  177 102 1·726±04 1·006±06   — — 
133GA VFR  441 198 1·364±10 1·014±10 1·174±12 0·861±10 — — 
61GC VFR  201  89·7 1·357±05 1·010±12 1·148±12 0·846±09  26·9±0·2 1·04±0·01 
55GC VFR  342 145 1·283±05 1·013±10 1·129±15 0·880±12 —  
L-1-4 VFR  398 176 1·344±04 1·007±06 1·185±18 0·881±13 — — 
L-1-5 VFR  400 177 1·346±06 0·992±05 1·183±10 0·878±07  91·3±3·5 1·73±0·07 
L-1-11 VFR  397 177 1·356±03 1·001±06 1·195±11 0·882±08 102·5±2·0 1·94±0·04 
L-1-83 VFR  408 180 1·340±06 1·005±08 1·207±29 0·901±22 — — 
L-2-5 VFR  352 159 1·368±04 0·996±04 1·179±16 0·862±12 103·2±2·7 1·88±0·07 
L-2-20 VFR  349 158 1·371±05 1·004±05 1·205±10 0·879±07 — — 
22-6-1 ELSC  155  51·0 1·000±03 1·004±05 1·109±16 1·110±16 — — 
20-5-2 ELSC   75·5  27·3 1·097±05 0·998±11 1·100±25 1·002±23 — — 
41-2-1 ILSC  108  34·6 0·977±04 1·007±06 1·077±09 1·102±10 — — 
10-1-3 CLSC  504 174 1·048±03 1·015±06 1·120±08 1·069±08  69·5±0·4 1·09±0·01 
11-2-1 CLSC  228  85·3 1·134±03 1·011±05 1·144±13 1·009±12  41·6±1·1 1·43±0·04 
13-2 CLSC 1123 390 1·054±04 1·013±04 1·128±14 1·070±13 142·1±1·6 1·01±0·02 
13-2 rpt CLSC 1123 392 1·059±04 1·023±09 1·104±06 1·043±07 — — 
12-5-2 CLSC  938 355 1·148±04 1·019±07 1·139±12 0·993±10 — — 
15-1-1 CLSC   90·0  32·9 1·111±08 1·001±08 1·203±28 1·083±26 — — 
46GC CLSC  103  35·1 1·030±03 1·033±13 1·101±09 1·070±09  25·2±0·4 1·96±0·04 
18GA2 CLSC  116  43·1 1·124±05 1·025±10 1·142±33 1·016±29 — — 
42GC CLSC   75·6  27·4 1·097±04 1·028±13 1·086±44 0·990±40  18·8±0·6 2·05±0·11 
Sample Location Th (238U/232Th) (234U/238U) (230Th/232Th) (230Th/238U) 226Ra (226Ra/230Th) 
 
 
(ppb)
 
(ppb)
 

 

 

 

 
(fg/g)
 

 
92KD1 VFR  349 156 1·355±06 1·011±08 1·213±12 0·895±10 — — 
90KD1 VFR  401 178 1·348±08 0·993±10 1·175±10 0·872±08 — — 
84KD1 VFR  138  61·3 1·347±04 1·019±11 1·170±08 0·869±07 — — 
114KD VFR  177 102 1·726±04 1·006±06   — — 
133GA VFR  441 198 1·364±10 1·014±10 1·174±12 0·861±10 — — 
61GC VFR  201  89·7 1·357±05 1·010±12 1·148±12 0·846±09  26·9±0·2 1·04±0·01 
55GC VFR  342 145 1·283±05 1·013±10 1·129±15 0·880±12 —  
L-1-4 VFR  398 176 1·344±04 1·007±06 1·185±18 0·881±13 — — 
L-1-5 VFR  400 177 1·346±06 0·992±05 1·183±10 0·878±07  91·3±3·5 1·73±0·07 
L-1-11 VFR  397 177 1·356±03 1·001±06 1·195±11 0·882±08 102·5±2·0 1·94±0·04 
L-1-83 VFR  408 180 1·340±06 1·005±08 1·207±29 0·901±22 — — 
L-2-5 VFR  352 159 1·368±04 0·996±04 1·179±16 0·862±12 103·2±2·7 1·88±0·07 
L-2-20 VFR  349 158 1·371±05 1·004±05 1·205±10 0·879±07 — — 
22-6-1 ELSC  155  51·0 1·000±03 1·004±05 1·109±16 1·110±16 — — 
20-5-2 ELSC   75·5  27·3 1·097±05 0·998±11 1·100±25 1·002±23 — — 
41-2-1 ILSC  108  34·6 0·977±04 1·007±06 1·077±09 1·102±10 — — 
10-1-3 CLSC  504 174 1·048±03 1·015±06 1·120±08 1·069±08  69·5±0·4 1·09±0·01 
11-2-1 CLSC  228  85·3 1·134±03 1·011±05 1·144±13 1·009±12  41·6±1·1 1·43±0·04 
13-2 CLSC 1123 390 1·054±04 1·013±04 1·128±14 1·070±13 142·1±1·6 1·01±0·02 
13-2 rpt CLSC 1123 392 1·059±04 1·023±09 1·104±06 1·043±07 — — 
12-5-2 CLSC  938 355 1·148±04 1·019±07 1·139±12 0·993±10 — — 
15-1-1 CLSC   90·0  32·9 1·111±08 1·001±08 1·203±28 1·083±26 — — 
46GC CLSC  103  35·1 1·030±03 1·033±13 1·101±09 1·070±09  25·2±0·4 1·96±0·04 
18GA2 CLSC  116  43·1 1·124±05 1·025±10 1·142±33 1·016±29 — — 
42GC CLSC   75·6  27·4 1·097±04 1·028±13 1·086±44 0·990±40  18·8±0·6 2·05±0·11 

Figure 2 also includes published U–Th analyses from the Tonga arc region for comparison (Regelous et al., 1997; Turner et al., 1997). Samples from the main Tonga arc (18°–21°S: Fig. 1) have very high (238U/232Th) of 2·2–3·2, and significant 238U excesses with (230Th/238U) of 0·55–0·80. Data from Ata island also plot to the right of the equiline, with (230Th/238U) of 0·81–0·95, and they have slightly higher (238U/232Th) than the adjacent Valu Fa Ridge (1·38–1·63 vs 1·28–1·37). The eruption ages for the Ata lavas are poorly constrained, and so the fact that samples with (238U/232Th) similar to the Valu Fa lavas show less 230Th–238U disequilibrium could simply be due to post-eruption decay (Ata lavas have 226Ra–230Th in equilibrium, implying an age ≫8 ka: Turner et al., 2000). The northern Tonga arc islands show less U enrichment than the main arc, with (238U/232Th) of 1·3–2·0, and (230Th/238U) of 0·69–0·95. The majority of samples analysed from the northern back-arc island of Niuafo’ou show 230Th excesses with (230Th/238U) of 0·94–1·41, at similar (238U/232Th) of 0·84–1·15 to the Central Lau Basin rift magmas (Regelous et al., 1997; Turner et al., 1997).

Fig. 2.

(a) (230Th/232Th) vs (238U/232Th) equiline diagram for the Tonga–Lau arc–back-arc system. •, samples from the central Tonga arc that have similar Nd isotopic compositions to the Valu Fa samples; together with the Valu Fa samples they define a linear trend with an apparent isochron age of 50 ka. Data sources: Tonga arc data and Niuafo’ou, Regelous et al. (1997) and Turner et al. (1997); Valu Fa, α-counting data, Vallier et al. (1991). (b) Close-up of part of the equiline diagram highlighting the new Lau Basin data. Of the nine samples analysed for 226Ra, eight have (226Ra/230Th) >1 (marked with × within the plotted symbols), and one has 226Ra/230Th in equilibrium (marked with a dot within the plotted symbol).

Fig. 2.

(a) (230Th/232Th) vs (238U/232Th) equiline diagram for the Tonga–Lau arc–back-arc system. •, samples from the central Tonga arc that have similar Nd isotopic compositions to the Valu Fa samples; together with the Valu Fa samples they define a linear trend with an apparent isochron age of 50 ka. Data sources: Tonga arc data and Niuafo’ou, Regelous et al. (1997) and Turner et al. (1997); Valu Fa, α-counting data, Vallier et al. (1991). (b) Close-up of part of the equiline diagram highlighting the new Lau Basin data. Of the nine samples analysed for 226Ra, eight have (226Ra/230Th) >1 (marked with × within the plotted symbols), and one has 226Ra/230Th in equilibrium (marked with a dot within the plotted symbol).

Valu Fa Ridge

Samples from the Valu Fa spreading centre have a relatively limited variation in both (238U/232Th) (1·34–1·37) and (230Th/232Th) (1·17–1·21), with the exception of the northernmost sample (55GC), which has a lower (238U/232Th) of 1·28 and (230Th/232Th) of 1·13. Sample 61GC, from an off-axis seamount, has similar (238U/232Th) to the adjacent spreading centre, but with lower (230Th/232Th) of 1·15. All samples show significant excess 238U with (230Th/238U) of 0·85–0·90. The L-1 and L-2 samples are from two closely spaced dredge hauls, and major and trace element data indicate that the samples within each dredge are compositionally very similar, perhaps fragments of the same flow (Vallier et al., 1991). This similarity is also apparent from the U-series results: (238U/232Th) and (230Th/232Th) are constant, within analytical error, at each dredge site, with the L-2 samples having slightly higher (238U/232Th) than those at L-1 (1·37 vs 1·35). These results provide a good estimate of the reproducibility of the data. The higher Th and U contents of the L-1 samples relative to those from L-2 (∼401 ppb vs ∼350 ppb, and ∼178 ppb vs ∼158 ppb, respectively) are consistent with their more fractionated major element compositions (see Table 1; Vallier et al., 1991). Most Valu Fa samples show significant 226Ra–230Th disequilibria (up to 100% excess 226Ra), confirming their relatively young eruption ages (<8 ka). The presence of young lava flows, many with evolved compositions, is consistent with the geophysical observations of a shallow magma chamber beneath the Valu Fa Ridge (Collier & Sinha, 1990).

Three of the samples (L-1-5, L-2-5, L-2-20) have previously been analysed for U-series nuclides by α-counting methods (Vallier et al., 1991). This earlier study similarly found significant excess 238U and excess 226Ra [(230Th/238U) 0·88–0·92, (226Ra/230Th) 1·4–2·6], and further showed that the samples were more than 100 years old because 210Pb and 226Ra were in secular equilibrium. However, there are significant discrepancies between these data and the new mass spectrometric data on the same samples, which are clearly illustrated in Fig. 2: the α-counting data show much higher (238U/232Th) and (230Th/232Th), 1·5 vs 1·2 and 1·7 vs 1·35, respectively. The U and Th concentrations determined by α-counting are also not consistent either with the new U and Th concentrations measured both by isotope dilution and by inductively coupled plasma mass spectrometry (ICP-MS) or with other trace element data on these samples (Tables 1 and 2).

Central Lau Basin (ELSC, ILSC and CLSC)

Samples from the Central Lau Basin have (238U/232Th) between 0·98 and 1·15, which is significantly lower than in Valu Fa Ridge samples (Fig. 2). All samples show either 230Th–238U equilibrium or 230Th excesses of up to 11%, in marked contrast to the Valu Fa Ridge samples from further south. The highest (230Th/238U) values are found in two samples from the ELSC and ILSC. Five samples have measured 230Th–238U equilibrium and although only two of these samples have been analysed for 226Ra, both show significant 226Ra–230Th disequilibrium, indicating eruption ages of <8 ka, and therefore that the 230Th–238U equilibrium is not due to post-eruption 230Th decay. One of the evolved andesitic samples (13-2) shows secular equilibrium for 226Ra–230Th. The maximum observed (226Ra/230Th) disequilibrium (∼2·0) is similar to that found at the Valu Fa Ridge.

Regional variation of 230Th–238U disequilibrium

Within the Tonga arc–back-arc system, there is a systematic difference in the sense of 230Th–238U disequilibrium away from the trench. This regional variation is summarized in Fig. 3, with (230Th/238U) plotted against distance behind the trench. It is clear, however, that there is not a progressive trend of 230Th–238U disequilibrium with distance, but rather an abrupt change in behaviour. Arc and back-arc lavas erupted <250 km from the trench have 238U excesses, whereas those from >250 km from the trench have 230Th excesses or 230Th–238U equilibrium. It should be noted that even though the CLSC is ∼100 km further from the trench than the ELSC, samples from both spreading centres have broadly similar extents of 230Th–238U disequilibrium, although the highest values are found closer to the arc in samples from the ELSC and ILSC.

Fig. 3.

(230Th/238U) vs distance behind the trench for samples from the Tonga arc and Lau Basin (data sources as for Fig. 2). The transition between lavas with 238U excesses and those with 230Th excesses occurs at ∼250 km from the trench.

Fig. 3.

(230Th/238U) vs distance behind the trench for samples from the Tonga arc and Lau Basin (data sources as for Fig. 2). The transition between lavas with 238U excesses and those with 230Th excesses occurs at ∼250 km from the trench.

DISCUSSION

Regional overview

A particular feature of the Tonga–Kermadec–Lau Basin system is the regional variation in the rates of back-arc spreading, in the compositions of mantle segments above the subduction zones, and in the components mobilized from the subducted oceanic crust. These provide exceptional opportunities to isolate contributions from well-characterized source materials, and to evaluate how the melt generation processes vary with distance from the trench, the time scales of the transfer of material from the subducted ocean crust, and whether there is any change in the hydrous fluid component with distance into the mantle wedge.

Identification of mantle wedge and slab-derived components in the Tonga region

Previous studies have established that the Tonga–Kermadec–Lau Basin system includes a number of compositionally distinct components and these need to be identified first. Most discussions of different components in the Tonga–Kermadec and Lau Basin rocks have relied on variations in Pb isotopes, and these are summarized in Fig. 4. Hergt & Hawkesworth (1994) documented two distinct trends in the Pb isotope data from rocks drilled at six sites during ODP Leg 135 in the older parts of the Lau Basin (Fig. 1). One trend includes samples from Sites 834 and 839 and projects back into the field for Pacific Ocean MORB (Fig. 4a), whereas the other trend includes samples from Sites 836 and 837 and projects into the field for Indian Ocean MORB. This was interpreted in terms of southward displacement of mantle similar to the source of Pacific MORB by Indian Ocean MORB mantle as a result of slab rollback, and accompanied by the southward migration of the propagating ridge tip into the extended crust (Hergt & Hawkesworth, 1994). As predicted by that model, the new data on the younger Central Lau Basin rocks have unradiogenic Pb isotope ratios that plot well within the field for Indian MORB (see also Loock et al., 1990), and the data for Ata plot in the array for Pacific MORB (Fig. 4a). The Valu Fa analyses plot close to where the two trends intercept, and so the significance of their Pb isotope ratios in the context of the material in the mantle wedge is more ambiguous. Bach et al. (1998) argued that instead of having two compositionally distinct mantle domains within the mantle wedge, the Pacific MORB isotopic signature can be explained simply by addition of fluids released by dehydration of the subducting Pacific plate. In the case of the Valu Fa Ridge, this fluid addition might swamp any contribution from the wedge, making it difficult to identify the isotopic provenance of the mantle wedge beneath this region.

By looking at the Pb isotope data in conjunction with trace element data, it is possible to distinguish between these two models, as well as to identify different components from the subducted slab. Ce/Pb is a particularly useful ratio in this regard. Ce and Pb have broadly similar incompatibility during mantle melting, and it has been argued that the mantle has a broadly constant Ce/Pb value of 25 ± 5 (Hofmann et al., 1986; Sims & DePaolo, 1997). In contrast, Pb is highly mobile in fluids relative to Ce (e.g. Miller et al., 1994). Arc rocks have relatively high Pb contents (and hence low Ce/Pb ratios), which are attributed to the preferential addition of Pb during subduction (e.g. Miller et al., 1994). The Pb isotope ratios of most arc rocks are therefore dominated by the contributions from subducted sediment and altered basalts, with little or no detectable contribution from the mantle wedge. Turner et al. (1997) used high Ba/Th ratios as a fingerprint of the fluid component in the Tonga arc lavas, and inferred that its 206Pb/204Pb ratio was ∼18·4, in contrast to the subducted sediments (low Ba/Th), which had 206Pb/204Pb >18·8. The variations in Ce/Pb ratios with 206Pb/204Pb are summarized in Fig. 4b, and it is important to note that binary mixing is linear in such a plot. The Central Lau Basin rocks scatter down to lower Ce/Pb ratios, consistent with a minor contribution from a Pb-rich fluid with 206Pb/204Pb = 18·4. The main array of the Tonga–Kermadec arc rocks reflects varying contributions from a similar fluid and a sediment component with 206Pb/204Pb >18·8. Evidence for the presence of mantle wedge material with a ‘Pacific’ MORB Pb isotope composition comes from some of the lavas from ODP Sites 834 and 839, which have high Ce/Pb values typical of MORB mantle unaffected by subduction fluids. The displacement of the Valu Fa Ridge rocks to higher Ce/Pb ratios relative to the main arc suggests less of a contribution from the subducted slab. It is consistent with the notion of the mantle wedge here having a ‘Pacific’ MORB composition similar to that tapped by lavas from ODP Sites 834 and 839, as opposed to the ‘Indian’ MORB mantle beneath the Central Lau Basin.

Fluid influences on melting and melt composition

A central issue is the amount of slab-released fluid present in the mantle where melting takes place, and whether the amount of fluid, particularly water, varies with the degree of melting and with distance from the trench. There are a few water analyses on Lau Basin rocks (Jambon & Zimmermann, 1990; Danushevsky et al., 1993; Kamenetsky et al., 1997; Bach et al., 1998; S. Newman, unpublished data, 1999), and these data indicate that the lavas are in general derived from mantle sources enriched in water relative to the average N-MORB source (e.g. Michael, 1995). In Fig. 5, enrichment in H2O relative to Ce (an element of similar incompatibility to water during mantle melting: Michael, 1995) is used to evaluate the variations in U/Th ratios. Ba/Th ratios are widely used to indicate the relative contribution of subduction-related fluids in arc rocks, but the plots of (238U/232Th) vs H2O/Ce and Ba/Th (Fig. 5) illustrate that two components may be identified in these rocks, in addition to MORB-type mantle in the wedge. One trend is to high (238U/232Th) with increasing Ba/Th and H2O/Ce (although the data for the latter are sparse) as defined by samples from Ata, which is the arc volcano furthest from the trench, and from Valu Fa, which is the back-arc section closest to the arc (Fig. 1). This is the generally accepted shift to high U/Th ratios attributed to the recent (<100 ky) introduction of fluids from the subducted crust. More surprisingly, the ELSC and CLSC rocks, which are further from the trench than those of the Valu Fa Ridge, define a second trend of slightly decreasing (238U/232Th) with increasing H2O/Ce accompanied by slight increases in Ba/Th (Fig. 5). They have higher H2O/Ce than average N-MORB (Michael, 1995), consistent with their more vesicular textures, and yet they are displaced to excess 230Th rather than excess 238U (Fig. 2). Thus, the water-rich component present in the mantle wedge in this region appears to be compositionally different from that found closer to the arc front. This could result from: (1) a similar slab-fluid, which had been compositionally modified through greater interaction with mantle wedge material (e.g. Stern et al., 1991; Hawkesworth et al., 1993; Stolper & Newman, 1994); (2) fluid dominantly derived from a different source, such as subducted sediments, which is inferred to have a longer residence time in the mantle wedge (e.g. Elliott et al., 1997; Turner & Hawkesworth, 1997; Class et al., 2000); and/or (3) it might be an inherent feature of the slightly trace element enriched Indian MORB source. These possibilities will be discussed in a later section.

Fig. 5.

(238U/232Th) vs (a) H2O/Ce, and (b) Ba/Th. H2O data by Fourier Transform Infrared spectroscopy (Jambon & Zimmermann, 1990; Kamenetsky et al., 1997; S. Newman, unpublished data, 1999). Average N-MORB data from GERM Web page, except average H2O/Ce (183 ± 33: Michael, 1995). Ata data from Turner et al. (1997). ◊, Seamount samples adjacent to the Valu Fa Ridge (Kamenetsky et al., 1997; this study). Highly evolved samples from the Central Lau Basin (grey shaded squares) have low Ba/Th, which is probably a result of extensive plagioclase fractionation (see Fig. 10b: Pearce et al., 1995). Trend 1, increasing U/Th with increasing H2O/Ce and Ba/Th, as shown by the Valu Fa Ridge and Ata samples near the trench; Trend 2, decreasing U/Th with increasing H2O/Ce and Ba/Th, as shown by the Central Lau Basin samples far from the trench.

Fig. 5.

(238U/232Th) vs (a) H2O/Ce, and (b) Ba/Th. H2O data by Fourier Transform Infrared spectroscopy (Jambon & Zimmermann, 1990; Kamenetsky et al., 1997; S. Newman, unpublished data, 1999). Average N-MORB data from GERM Web page, except average H2O/Ce (183 ± 33: Michael, 1995). Ata data from Turner et al. (1997). ◊, Seamount samples adjacent to the Valu Fa Ridge (Kamenetsky et al., 1997; this study). Highly evolved samples from the Central Lau Basin (grey shaded squares) have low Ba/Th, which is probably a result of extensive plagioclase fractionation (see Fig. 10b: Pearce et al., 1995). Trend 1, increasing U/Th with increasing H2O/Ce and Ba/Th, as shown by the Valu Fa Ridge and Ata samples near the trench; Trend 2, decreasing U/Th with increasing H2O/Ce and Ba/Th, as shown by the Central Lau Basin samples far from the trench.

Stolper & Newman (1994) reported a good positive correlation between H2O and U contents in glasses from the Mariana Trough, and a similar correlation is observed between U and H2O for the Lau Basin samples for which water data are available (Fig. 6a). This indicates that, to a first approximation, U can be used as a proxy for H2O, although we note that some H2O is likely to have been lost by degassing particularly in the evolved Valu Fa Ridge lavas and and that subduction fluids might have somewhat variable U/H2O values. Yb is plotted against MgO for the rocks of the Lau Basin and the Tonga arc in Fig. 6b, to emphasize that most suites include rocks with ∼8% MgO and that there are significant differences in the Yb contents of the less evolved rocks. The plot of U vs Yb (Fig. 6c) shows a series of steep arrays for the different rock suites, primarily reflecting within-suite differentiation. However, the low ends of these arrays tend to be in rocks with ∼8% MgO (Fig. 6b), and in those least evolved rocks there is a progressive shift to decreasing Yb contents with increasing U. As U reflects the fluid contribution, and may be a rough proxy for H2O in these rocks, and Yb varies with the degree of melting (see also Fig. 8, below), the arrays in Fig. 6c strongly suggest that the degree of melting decreases with decreasing fluid (water) contribution from the subducted slab with distance away from the trench. In general, however, within arc rocks, the relative contribution from the fluid component tends to be high (e.g. high Ba/Th, U/Th) in rocks derived from more depleted sources (e.g. high Al/Ti, low Na/Ta) (e.g. Ewart & Hawkesworth, 1987; Hawkesworth et al., 1991, 1997b; Woodhead et al., 1993; Turner et al., 1997), but there is less evidence that the degree of melting varies solely with the amount of added fluid. Global inter-arc correlations of major elements with lithospheric thickness (Plank & Langmuir, 1988), theoretical modelling of melting of hydrated peridotite (Hirschmann et al., 1999), differences in fO2 between primitive arc lavas and arc peridotites (Parkinson & Arculus, 1999), and the existence of primitive low-H2O basalts in some arc-front volcanoes (Sisson & Bronto, 1998) all suggest a component of decompression melting beneath arcs. Thus, the degree of melting beneath arcs is likely to be largely determined by a combination of volatile addition and decompression melting (e.g. Pearce & Parkinson, 1993; Pearce & Peate, 1995).

Fig. 8.

Nb vs Yb diagram for primitive lavas from the Lau Basin–Tonga arc system, to resolve partial melting trends from source depletion trends (Pearce & Parkinson, 1993). All plotted samples have 7–8 wt % MgO, except the Central Tonga arc (6–7 wt % MgO). It should be noted that the grid is calibrated for samples with 9 wt % MgO (Pearce & Parkinson, 1993). FMM, fertile MORB mantle, equivalent to the N-MORB source.

Fig. 8.

Nb vs Yb diagram for primitive lavas from the Lau Basin–Tonga arc system, to resolve partial melting trends from source depletion trends (Pearce & Parkinson, 1993). All plotted samples have 7–8 wt % MgO, except the Central Tonga arc (6–7 wt % MgO). It should be noted that the grid is calibrated for samples with 9 wt % MgO (Pearce & Parkinson, 1993). FMM, fertile MORB mantle, equivalent to the N-MORB source.

The Valu Fa Ridge

Fluid transfer rates at the Valu Fa Ridge

The Valu Fa Ridge lies in close proximity to the Tonga arc front (40 km from Ata) and it is only 150 km above the Benioff zone. Previous studies (e.g. Vallier et al., 1991; Bach et al., 1998) have shown that the Valu Fa magmas have pronounced arc-like compositions. This is illustrated both in the MORB-normalized trace element plot in Fig. 7 (e.g. Pearce & Peate, 1995), where the contrast to the MORB-like samples from the CLSC and ELSC is clear, and by the higher U/Th of the Valu Fa magmas seen in Fig. 2. At issue is the origin of this strong subduction influence in a back-arc rift. Boespflug et al. (1990) suggested that this may be a result of the migration of the arc volcanic front into the back-arc region. The Valu Fa Ridge is propagating southwards into older crust, but the age and composition of that crust are essentially unknown (Hawkins, 1995a). Most workers have suggested that the Valu Fa magmas tap a depleted MORB mantle source plus a slab-derived component (Jenner et al., 1987; Boespflug et al., 1990; Loock et al., 1990; Vallier et al., 1991; Bach & Niedermann, 1998; Bach et al., 1998), but some have also raised the possibility that some of the subduction characteristics are acquired through interaction with this older crust during rifting (Vallier et al., 1991; Hilton et al., 1993; Hawkins, 1995a). Hilton et al. (1993) showed that 3He/4He of Valu Fa magmas is negatively correlated with SiO2 and that the more evolved samples have extremely low 3He/4He (R/RA ∼1): in detail, samples with 53–55% SiO2 had 6–8 R/RA, and it was only samples with >58% SiO2 that had <3 R/RA. Hilton et al. suggested that these He results could be explained by assimilation of Lau crust by previously degassed magmas, and this model requires that parts of this basement crust are old enough to have in-grown sufficient radiogenic 4He. However, such shallow-level assimilation does not appear to have markedly influenced the U–Th data. This can be seen from comparing the most primitive (84KD, SiO2 51%) and most evolved (133GA, SiO2 57%) samples from the Valu Fa spreading centre: both samples have identical (238U/232Th) and (230Th/232Th), despite a three-fold enrichment in Th content. Furthermore, the significant 238U–230Th disequilibrium in all Valu Fa magmas is not consistent with any model in which the subduction characteristics of the Valu Fa magmas are derived from shallow-level bulk assimilation of arc-related crust, as this crust would be expected to be old enough to have 238U–230Th in equilibrium.

Fig. 7.

N-MORB-normalized trace element patterns for representative samples from the Lau Basin region, all with similar MgO contents. CLSC (15-1-1: MgO 7·3 wt %), ELSC (22-6-1: MgO 7·1 wt %), VFR (84KD1: MgO 7·3 wt %), and Ata (482-8-11: MgO 7·4 wt %, from Turner et al., 1997). Normalizing values from GERM Web page.

Fig. 7.

N-MORB-normalized trace element patterns for representative samples from the Lau Basin region, all with similar MgO contents. CLSC (15-1-1: MgO 7·3 wt %), ELSC (22-6-1: MgO 7·1 wt %), VFR (84KD1: MgO 7·3 wt %), and Ata (482-8-11: MgO 7·4 wt %, from Turner et al., 1997). Normalizing values from GERM Web page.

Excesses of 238U over 230Th that are characteristic of many arc rocks are generally attributed to the recent addition of a U-rich fluid (e.g. Newman et al., 1984; Gill & Williams, 1990; Sigmarsson et al., 1990; McDermott & Hawkesworth, 1991; Condomines & Sigmarsson, 1993; Elliott et al., 1997; Hawkesworth et al., 1997a; Turner et al., 1997). In detail, suites of well-dated samples from individual arcs often define linear trends on an equiline diagram, which, if interpreted as isochrons, yield ages of between 20 and 70 ka (e.g. Sigmarsson et al., 1990; Elliott et al., 1997; Hawkesworth et al., 1997a; Turner et al., 1997, 1999). The U/Th ratios, and hence the arrays on the U–Th equiline diagrams, are interpreted as mixtures between wedge + sediment (low U/Th) and fluid from the subducted slab (high U/Th). As the fluid is inferred to contain little or no Th (e.g. Hawkesworth et al., 1997b), the mixing arrays are horizontal at the time of fluid addition, and the slope after subsequent 230Th in-growth is inferred to indicate the age since fluid addition. The low-U/Th end-member may be regionally variable as a result of differences in composition of mantle wedge and proportions of added sediment material. Some of the scatter in the 238U–230Th data from the Tonga arc reflects variable sediment contributions, and the scatter is significantly reduced by considering only those samples with similar Nd isotope ratios, as Nd is not thought to be present in significant quantities in the fluid component (e.g. Hawkesworth et al., 1997b; see below).

Lavas from the Central Tonga arc all have high (238U/232Th) and significant 230Th–238U disequilibrium (Regelous et al., 1997; Turner et al., 1997), but it is difficult to constrain the slope of an array through the data, because of the relatively restricted range in U/Th. The northern Tonga lavas have lower U/Th (Fig. 2), but their relationship to the Central Tonga lavas is complicated by the presence of additional components in their mantle source (Regelous et al., 1997; Turner & Hawkesworth, 1997, 1998; Turner et al., 1997; Wendt et al., 1997; Ewart et al., 1998). Lavas from Ata also have lower U/Th, but their age is uncertain so that the data may require significant age corrections on the equiline diagram. The Valu Fa magmas are another potential low-U/Th candidate, and it is therefore important to assess their relationship to the arc-front magmatism.

Pearce & Parkinson (1993) demonstrated that a plot of Nb vs Yb, using analyses of primitive basalts, can be effective in resolving partial melting differences from source depletion effects. Valu Fa basalts fall on the same trend as that defined by lavas from Ata and most of the Central Tonga arc (except Tofua), and this trend can be explained by different degrees of melting of a similar, depleted source (Fig. 8). Thus, it seems reasonable to infer that the mantle source at the Valu Fa Ridge is broadly similar to that for the Central Tonga arc-front volcanoes and that the Valu Fa magmas represent the best estimate of the low-U/Th, fluid-poor, end-member. Therefore, they can be used to anchor a trend line through data from the Central Tonga arc on the equiline diagram. To minimize the effects of variable sediment addition, the Central Tonga arc data are restricted to samples with similar Nd isotope compositions to the Valu Fa magmas (143Nd/144Nd 0·51302–0·51307: Turner et al., 1997; Bach et al., 1998; Ewart et al., 1998). The resulting trend line (Fig. 2) can be interpreted as an isochron, indicating an age of ∼50 ka. This is within error of the age inferred from 231Pa–235U systematics of samples from just the Tonga arc (Bourdon et al., 1999). Both trends can be most simply explained by the addition of a U-rich, slab-derived fluid to the mantle source at ∼50–60 ka. Thus, the time scales of fluid transfer from the subducted slab appear to be similar beneath the Valu Fa back-arc ridge and beneath the arc itself. This is consistent with a tectonic model in which the Valu Fa Ridge is propagating into the arc-front region.

Melt generation processes at the Valu Fa Ridge

Previous studies have shown that the Valu Fa basalts sample a refractory mantle wedge, as indicated by the presence of Mg-rich olivines and Cr-rich spinels (Sunkel, 1990; Kamenetsky et al., 1997; Bach et al., 1998). High field strength element ratios such as Zr/Nb and Ti/Zr indicate that this mantle is more depleted than an N-MORB source. A recent study on melt inclusions in olivines from near-axis seamount lavas on both sides of the Valu Fa Ridge spreading centre has highlighted the complexities of melt generation processes and the mineralogy and trace element composition of the diverse mantle sources beneath the Valu Fa Ridge–Ata region (Kamenetsky et al., 1997). Major element compositions of the melt inclusions suggest a refractory, harzburgitic, hydrated sub-arc lithosphere source for the seamount lavas, variably veined with clinopyroxene-rich dykes closer to the arc. Trace element data on the melt inclusions, coupled with whole-rock analyses of lavas from the Valu Fa Ridge itself and Ata, indicate the presence of three distinct components within this region. The seamount lavas and their melt inclusions, plus Ata, fall between two components on many trace element plots: a fluid-rich component with high Ba/Th and low Ce/Pb, dominant in the eastern seamounts, and a ‘boninitic’ component with high La/Yb and low Ba/Th, dominant in the western seamounts. The Valu Fa Ridge samples are displaced from these trends towards unmodified MORB-type mantle wedge compositions.

The seamount magmas have been interpreted as the products of melting of shallow, hydrated sub-arc lithosphere as a result of conductive heating and decompression caused by entrainment into upwelling MORB-source mantle of the developing Valu Fa Ridge (Kamenetsky et al., 1997). In this model, the decrease in ‘subduction’ signature from east to west would be due to a decrease in the extent of ‘fluid’ addition to the lithosphere before rifting. However, conductive heating is likely to take longer than fluid-induced melting, which in turn suggests that any excess 238U as a result of fluid addition will have decayed back to isotope equilibrium. Only one seamount sample has been analysed (61GC), and this has excess 238U similar to the other Valu Fa rocks, consistent with fluid-induced melting this close to the arc front. Although it is likely that hydrated asthenosphere can decompress (and melt) to shallower depths beneath the Valu Fa spreading centre than beneath the Ata volcano, the degree of melting is apparently higher at Ata than at the Valu Fa Ridge, as is indicated by Fig. 8. Figure 6 suggests that the additional melting beneath Ata can be linked to a higher fluid flux.

Fig. 6.

(a) U vs H2O, (b) Yb vs MgO, (c) U vs Yb. In (a), the outlined field highlights the more primitive (7–8 wt % MgO) samples. The solid arrows indicate a trend of inferred decreasing fluid content and decreasing degree of melting in primitive arc and back-arc lavas. The dashed arrows represent schematic within-suite crystal fractionation trends. Data sources as for Figs 4 and 5.

Fig. 6.

(a) U vs H2O, (b) Yb vs MgO, (c) U vs Yb. In (a), the outlined field highlights the more primitive (7–8 wt % MgO) samples. The solid arrows indicate a trend of inferred decreasing fluid content and decreasing degree of melting in primitive arc and back-arc lavas. The dashed arrows represent schematic within-suite crystal fractionation trends. Data sources as for Figs 4 and 5.

The Central Lau Basin

Characterization of mantle sources beneath the Central Lau Basin

The Central Lau Basin spreading centres are further behind the trench than the Valu Fa Ridge, and the subduction influence on magma compositions is correspondingly diminished (e.g. Hawkins, 1995a; Pearce et al., 1995). All the lavas are broadly N-MORB in composition, as can be seen in the MORB-normalized trace element diagram (Fig. 7). It is important to clarify whether the CLSC lavas have any compositional features to distinguish them from N-MORB, which might indicate whether any slab-derived fluid related to active subduction has reached that far behind the arc. Major element trends shown by CLSC lavas are similar to MORB (Hawkins, 1995a; Pearce et al., 1995), but there are significant minor differences in selected trace element ratios relative to average N-MORB values: e.g. primitive CLSC basalts have higher Ba/Th, Ba/Nb and K/Nb, and lower Ce/Pb (Figs 4, 5 and 9). The highly fractionated CLSC lavas are excluded from these discussions because they have low Ba/Th and high Ce/Pb as a result of extensive plagioclase fractionation (Pearce et al., 1995): both ratios show good correlations with Eu/Eu* (e.g. Fig. 10b). These trace element characteristics indicate the influence of a subduction-related component, and at first sight it is tempting to attribute them to the muted effect of material added by contemporaneous subduction processes. However, it is important to note that average N-MORB values (e.g. GERM Web page) are defined based primarily on samples of Atlantic and Pacific MORB. The Pb isotope data summarized earlier clearly show that the CLSC lavas are derived from a source similar to that of Indian MORB (Loock et al., 1990; Hergt & Hawkesworth, 1994). Recent studies have indicated that the Indian MORB source is distinct from the Atlantic and Pacific MORB sources not only in terms of isotope composition but in trace element composition as well, with Indian MORB characterized by enrichments in Ba and Pb and depletions in Nb relative to Atlantic and Pacific MORB (e.g. Rehkämper & Hofmann, 1997). Figure 9 indicates that the CLSC basalts have Ba/Nb ratios and Sr isotope compositions that overlap with the more enriched end of the Indian MORB compositional array. In terms of volatiles, the CLSC lavas are enriched in H2O (H2O/Ce = 345 ± 69 for CLSC (Jambon & Zimmermann, 1990; Danushevsky et al., 1993; S. Newman, unpublished data, 1999) relative to average N-MORB (183 ± 33 for N-MORB: Michael, 1995) and although water data are sparse on Indian MORB samples, preliminary data suggest that the Indian MORB source might also be characterized by relatively high H2O/Ce (Michael, 1995). Two CLSC samples measured for Li isotopes have higher δ6Li (−0·5‰ and −1·7‰) than the apparent range for Atlantic and Pacific MORB (−3·4‰ to −4·7‰) (Chan et al., 1999), but no data are yet available to verify whether this is also a feature of the Indian MORB source. Major element systematics also indicate that the primitive CLSC lavas were formed by melting of an Indian MORB mantle source through processes typical of a mid-ocean ridge spreading centre. Basaltic lavas from the CLSC plot on the broad correlations found between axial depth, Na8·0 and Fe8·0 in MORB globally (the subscript denotes that the composition is fractionation corrected to an MgO of 8 wt %: Klein & Langmuir, 1987). Furthermore, in detail, Indian MORB basalts have systematically lower Fe8·0 than Atlantic MORB basalts at a given depth (Langmuir et al., 1992), and this is also shown by the CLSC basalts (data in Pearce et al., 1995). Thus, it appears that the distinctive compositional features of the CLSC mantle source owe their origin primarily to ancient enrichments that are characteristic of the Indian MORB mantle source rather than being related to contemporaneous subduction processes (see also Pearce et al., 1995).

Fig. 9.

Ba/Nb vs 87Sr/86Sr to illustrate the compositional similarity between Central Lau Basin basalts (CLSC, ELSC, ILSC) and some Indian MORB. Only the primitive Central Lau Basin samples (7–8 wt % MgO) are plotted. MORB data from Rehkämper & Hofmann (1997), and references therein, plus the GERM Web page.

Fig. 9.

Ba/Nb vs 87Sr/86Sr to illustrate the compositional similarity between Central Lau Basin basalts (CLSC, ELSC, ILSC) and some Indian MORB. Only the primitive Central Lau Basin samples (7–8 wt % MgO) are plotted. MORB data from Rehkämper & Hofmann (1997), and references therein, plus the GERM Web page.

The evolved rocks of the CLSC have undergone extensive crystallization of a plagioclase-dominated assemblage from parental basaltic magmas as a result of a high rate of cooling relative to fresh magma supply associated with the propagating rift tip (Pearce et al., 1995). Oxygen isotope data suggest that the precursor magmas to these highly fractionated lavas had assimilated hydrothermally altered oceanic crust material (Macpherson & Mattey, 1998). The slightly elevated 87Sr/86Sr but similar 143Nd/144Nd of the evolved rocks relative to the basalts on the CLSC (Fig. 10a) is consistent with such a model. The few Ra data available show a positive correlation between (226Ra/230Th) and Eu/Eu* (Fig. 10c), which could be explained by extensive plagioclase fractionation (see also Turner et al., 2000). However, it is difficult to evaluate the effects of radioactive decay for the evolved samples in the absence of independent eruption age information. There is no clear relationship, though, between (238U/232Th) or (230Th/238U) with either a monitor of crystallization such as Eu/Eu* or a monitor of assimilation such as 87Sr/86Sr that might otherwise indicate an important influence on the 238U–230Th disequilibria data.

The ELSC and ILSC lie closer to the arc than the CLSC, and so any compositional differences between lavas from these two areas might reflect processes associated with the present subduction setting. Major element features of the ELSC lavas, such as lower Na2O and Fe2O3 and higher SiO2 at a given MgO, coupled with an earlier onset of oxide and apatite saturation, all indicate that the mantle beneath the ELSC is, in general, more hydrous and more depleted than the CLSC mantle (Hawkins, 1995a; Pearce et al., 1995). This is consistent with the observation that ELSC and ILSC glasses have much higher H2O/Ce (788–1266) than CLSC samples (224–470) (Jambon & Zimmermann, 1990; Danushevsky et al., 1993; S. Newman, unpublished data, 1999). The ELSC and ILSC basalts included in this study have slightly higher Ba/Th (Fig. 5) and higher 206Pb/204Pb (Fig. 4) than the CLSC basalts. ELSC samples dredged just to the south of the studied lavas are more highly vesicular, despite similar ridge depths, which suggests an even greater volatile content, and they also have higher Ba/Th (Pearce et al., 1995). The elevated H2O/Ce ratios for the ELSC and ILSC relative to the CLSC are also much greater than found in any MORB globally (<380: Michael, 1995) and the fact that these increase with arc proximity, along with Ba/Th, suggests that this reflects an additional fluid component related to the active subduction environment, rather than being a more extreme variety of Indian-MORB-type mantle (Fig. 5).

Fig. 4.

206Pb/204Pb vs (a) 208Pb/204Pb, and (b) Ce/Pb, to illustrate the different mantle and slab components involved in magmatism in the Tonga–Kermadec arc and the Lau Basin. Data sources: Jenner et al. (1987), Vallier et al. (1991), Ewart et al. (1994, 1998), Hergt & Farley (1994), Regelous et al. (1997), Turner et al. (1997), Bach et al. (1998) and this paper. ‘Indian’ MORB mantle has lower Ce/Pb than ‘Pacific’ MORB mantle (Rehkämper & Hofmann, 1997; Sims & DePaolo, 1997). Circle with cross represents average ‘Indian’ MORB from Rehkämper & Hofmann (1997). Highly evolved samples from the Central Lau Basin (grey shaded squares: MgO ≪5 wt%) have elevated Ce/Pb (>20), which is probably a result of extensive plagioclase fractionation (marked by dashed arrow), e.g. Sims & DePaolo (1997).

Fig. 4.

206Pb/204Pb vs (a) 208Pb/204Pb, and (b) Ce/Pb, to illustrate the different mantle and slab components involved in magmatism in the Tonga–Kermadec arc and the Lau Basin. Data sources: Jenner et al. (1987), Vallier et al. (1991), Ewart et al. (1994, 1998), Hergt & Farley (1994), Regelous et al. (1997), Turner et al. (1997), Bach et al. (1998) and this paper. ‘Indian’ MORB mantle has lower Ce/Pb than ‘Pacific’ MORB mantle (Rehkämper & Hofmann, 1997; Sims & DePaolo, 1997). Circle with cross represents average ‘Indian’ MORB from Rehkämper & Hofmann (1997). Highly evolved samples from the Central Lau Basin (grey shaded squares: MgO ≪5 wt%) have elevated Ce/Pb (>20), which is probably a result of extensive plagioclase fractionation (marked by dashed arrow), e.g. Sims & DePaolo (1997).

In the northern ELSC and ILSC, only increased water contents are detectable, whereas closer to the arc, south of 20°S, Ba abundances are also clearly elevated, indicating that the composition of the subduction component varies systematically towards the arc, as suggested by Pearce et al. (1995). Furthermore, the fluid component in the ELSC and ILSC mantle source is apparently different in composition from that found near the arc front, as discussed above, and one of the principal differences is that it does not have a high U/Th ratio (see Fig. 5). One plausible model to account for this regional variation in fluid composition is a chromatographic model (e.g. Stern et al., 1991; Hawkesworth et al., 1993; Stolper & Newman, 1994), in which the slab-derived fluid progressively equilibrates with mantle material during transport through the wedge. The fluid gradually loses its subduction signature depending on the length of the fluid path through the mantle and the fluid–mantle partition coefficients: only those elements with low fluid–mantle partition coefficients (e.g. Ba, K) will retain evidence of their slab origin at significant distances behind the trench. The calculated subduction fluid composition for the Mariana Trough back-arc basalts also had a relatively low (238U/232Th) of ∼1·19 ± 0·07 (Stolper & Newman, 1994).

Melt generation processes in the Central Lau Basin

The Central Lau Basin lavas are all characterized by having (230Th/238U) ≥1, which clearly distinguishes them from the arc-like Valu Fa samples with (230Th/238U) <1 (Fig. 2). The variations shown by the Central Lau Basin samples on the equiline diagram cannot be explained by addition of a high-U/Th fluid similar to that at Valu Fa to a melt with low U/Th and high (230Th/238U), as it is the two ‘fluid-rich’ ELSC and ILSC samples that have the lowest U/Th and highest (230Th/238U). Instead, the behaviour of U and Th in melts generated at these distal back-arc spreading centres appears to be governed by processes common to typical mid-ocean ridges. Back-arc magmas from the CLSC, where there is no detectable subduction contribution, primarily show 230Th excesses, as also seen on the back-arc island of Niuafo’ou (Figs 13: Regelous et al., 1997; Turner et al., 1997), indicating generation by decompression melting initiated within the garnet stability field. The more water-rich magmas from the ELSC and ILSC tend to have even greater 230Th excesses.

The Central Lau Basin data form a broad positive array on the equiline diagram (Fig. 2), reminiscent of the trends shown by MORB from individual ridge segments globally (Lundstrom et al., 1998b). Such trends have been explained in terms of mixing between melts derived from a compositionally heterogeneous source, with ‘enriched’ sources having lower (238U/232Th) than ‘depleted’ sources, and their slopes appear to be largely dependent on the local spreading rate (Lundstrom et al., 1998b). A trend through the Central Lau Basin data has a calculated slope of 0·45 (0·85 if 15-1-1 is included). These values are broadly consistent with the range in model-dependent values of 0·4–0·7 predicted from Lundstrom et al. (1998b), assuming a half-spreading rate of ∼45 mm/yr, which is reasonable for both CLSC and ELSC (Taylor et al., 1996; Zellmer & Taylor, 1999). The time-integrated U/Th ratio of the mantle beneath the Central Lau Basin can be estimated from the Pb isotope composition of these lavas using the radiogenic 208Pb*/206Pb*–(230Th/232Th) global mantle array (Allègre et al., 1986). A 208Pb*/206Pb* value of ∼0·959 for the Central Lau Basin lavas gives a (230Th/232Th) of ∼1·04, which corresponds to the mantle source (238U/232Th) at equilibrium. This is similar to the value of ∼1·1 inferred for the Indian MORB source beneath the Vanuatu arc from Pb isotopes and U-series disequilibrium data (Turner et al., 1999). For ridge settings, it has been argued that the measured sample U/Th provides a better estimate of source U/Th than the (230Th/232Th) ratio (e.g. O’Nions & McKenzie, 1993; Elliott, 1997). The measured range in (238U/232Th) in the Central Lau Basin glasses is 0·98–1·15, which is also consistent with the value estimated from Pb isotopes.

Global MORB display a broad negative correlation between (230Th/238U) and axial ridge depth (Fig. 11) (Bourdon et al., 1996b). The correlation is improved if averages for ridge segments are plotted instead, because there can be a wide range in (230Th/238U) within a particular area and depth range as a result of local source heterogeneities (Bourdon et al., 1996b; Lundstrom et al., 1998b). Global correlations between major element compositions of MORB and axial depth have been explained by mantle temperature variations (Klein & Langmuir, 1987; Langmuir et al., 1992), and this model can also account for the global trend between 230Th–238U disequilibrium and depth (Bourdon et al., 1996b). The Central Lau Basin data plot below the broad global MORB array in Fig. 11. They have lower values of (230Th/238U) than expected for their relatively shallow water depths. Such relatively low 230Th excesses could be explained by a faster mantle upwelling rate beneath these spreading centres (Bourdon et al., 1996b), which might be related to the broader mantle dynamics within the wedge, where ‘Indian’-type mantle is actively displacing ‘Pacific’-type mantle (Hergt & Hawkesworth, 1994). It is difficult independently to assess the upwelling rate to confirm such a model, although 231Pa–235U data in conjunction with the 230Th–238U results might help to better define plausible estimates (e.g. Lundstrom et al., 1998a). It is also possible that the relationship between axial depth and mantle temperature is different between mid-ocean ridges and back-arc basins, with the tectonic environment having a regional effect on back-arc basin depths. Park et al. (1991) noted that for young back-arc basins (<10 Ma), the axial depth is correlated with the dipping angle of the slab: as the dip increases, the basin depth becomes larger.

Fig. 10.

(a) 87Sr/86Sr vs 143Nd/144Nd. The slightly elevated 87Sr/86Sr of evolved CLSC lavas compared with the basalts suggests minor assimilation of altered oceanic crust. (b) Eu/Eu* vs Ba/Th. Whereas Ba/Th is high in altered oceanic crust (∼323: Staudigel et al., 1996), the lower Ba/Th in the evolved CLSC lavas compared with the basalts indicates that the effects of extensive plagioclase fractionation dominate over any minor assimilation of altered oceanic crust. (c) Eu/Eu* vs (226Ra/230Th). Although plagioclase crystallization should fractionate 226Ra/230Th in a similar manner to Ba/Th, in the absence of independent eruption age estimates, the effects of radioactive decay cannot be assessed.

Fig. 10.

(a) 87Sr/86Sr vs 143Nd/144Nd. The slightly elevated 87Sr/86Sr of evolved CLSC lavas compared with the basalts suggests minor assimilation of altered oceanic crust. (b) Eu/Eu* vs Ba/Th. Whereas Ba/Th is high in altered oceanic crust (∼323: Staudigel et al., 1996), the lower Ba/Th in the evolved CLSC lavas compared with the basalts indicates that the effects of extensive plagioclase fractionation dominate over any minor assimilation of altered oceanic crust. (c) Eu/Eu* vs (226Ra/230Th). Although plagioclase crystallization should fractionate 226Ra/230Th in a similar manner to Ba/Th, in the absence of independent eruption age estimates, the effects of radioactive decay cannot be assessed.

Fig. 11.

(230Th/238U) vs axial ridge depth. Global MORB data by TIMS from Goldstein et al. (1991), Volpe & Goldstein (1993), Lundstrom et al. (1995), Bourdon et al. (1996a, 1996b), Lundstrom et al. (1998a, 1998b), Peate et al. (2001). Mariana Trough data are α-counting analyses from Newman (1983).

Fig. 11.

(230Th/238U) vs axial ridge depth. Global MORB data by TIMS from Goldstein et al. (1991), Volpe & Goldstein (1993), Lundstrom et al. (1995), Bourdon et al. (1996a, 1996b), Lundstrom et al. (1998a, 1998b), Peate et al. (2001). Mariana Trough data are α-counting analyses from Newman (1983).

Fig. 12.

Schematic cross-sections [modified from Hawkins (1995a)] through the Lau Basin–Tonga arc system, at; (a) ∼19°S, (b) ∼22°S. The two sections are aligned on the arc front, and the inset shows how U/Th decreases laterally away from the trench.

Fig. 12.

Schematic cross-sections [modified from Hawkins (1995a)] through the Lau Basin–Tonga arc system, at; (a) ∼19°S, (b) ∼22°S. The two sections are aligned on the arc front, and the inset shows how U/Th decreases laterally away from the trench.

An additional complication in the back-arc environment is the effect of an elevated and variable water content of the mantle source during melting, which is likely to have a significant influence on the systematics of 230Th–238U but not in an easily quantifiable way yet. An increased water content will lower the mantle solidus temperature and cause melting to be initiated at a greater depth. This leads to a greater garnet influence, which produces greater 230Th excesses for an ‘in-growth’ model (e.g. Bourdon et al., 1996b; Lundstrom et al., 1998b). Thus, this effect of water would be similar to that caused by an elevated mantle temperature, which also increases the depth and extent of melting and results in the high 230Th excesses observed in MORB from shallow ridges (Fig. 11: Bourdon et al., 1996b; Lundstrom et al., 1998b). However, another important consequence of an elevated water content is its effect on melt productivity. It appears that melt productivity increases with increasing water content, which will reduce the amount of 230Th in-growth during melting (Lundstrom et al., 1998b; Hirschmann et al., 1999). The subdued 230Th excesses of the Central Lau Basin glasses suggest that the influence of water on melt productivity might be more important in governing the systematics of U-series disequilibrium in these melts than its effect on lowering the mantle solidus temperature.

The behaviour of uranium is strongly dependent on the prevailing redox environment. Arc peridotite data indicate that conditions in the sub-arc mantle are markedly more oxidizing than those in oceanic or ancient cratonic mantle as a result of the influx of slab-derived fluids, and that the sub-arc mantle is oxidizing enough for U6+ to be the dominant uranium species, both in the source region and during melting (Parkinson & Arculus, 1999). In Central Lau Basin basaltic glasses, fO2 is positively correlated with H2O content, which indicates that relatively oxidized mantle exists this far behind the arc (Farley & Newman, 1994). The highly incompatible nature of U and Th means that 230Th–238U disequilibrium is strongly influenced by processes within the initial melting regime (<1% melt) at the base of the melting column. However, the partitioning behaviour of uranium during these initial stages of melting, when mantle peridotite will be in equilibrium with oxidizing, water- and alkali-rich melts, is poorly constrained. It has been suggested that U would be more incompatible than Th in such oxidizing conditions, thus resulting in reduced 230Th excesses (Bourdon et al., 1996b) as is observed in the Central Lau glasses as a whole. However, in detail, the ELSC–ILSC magmas are more oxidizing than the CLSC magmas (Pearce et al., 1995) and yet they tend to have higher (230Th/238U).

Back-arcs in general show similar global systematics to other ocean ridges in terms of correlations between Fe8·0, Na8·0 and axial depth, but they are offset systematically to lower Fe8·0 (Langmuir et al., 1992). The relative paucity of information about the effects of water on basalt phase equilibria and during mantle melting means that this difference has not been quantitatively modelled yet. Details of the Na8·0–Fe8·0–depth relationships of the Central Lau Basin basalts have been discussed by Pearce et al. (1995). In a diagram of Na8·0 vs Fe8·0, ELSC magmas fall on a so-called ‘local’ trend of decreasing Fe8·0 with decreasing Na8·0, characteristic of slow-spreading mid-ocean ridges (Klein & Langmuir, 1987; Langmuir et al., 1992). This ‘local’ trend for the ELSC intersects the broad negative array defined by global MORB at lower Na8·0 than that of the CLSC samples. This suggests a higher degree of melting for ELSC magmas relative to CLSC magmas, which is consistent with their higher water contents (Stolper & Newman, 1994; Pearce et al., 1995). Samples from closer to the arc along the ELSC plot further from the global trend to low Na8·0 and low Fe8·0, as does the sample from the ILSC (41-2-1). Thus, the two samples with the lowest U/Th and highest (230Th/238U) in the Central Lau Basin (41-2-1 and 22-6-1) plot at either end of the Na8·0–Fe8·0 ‘local’ trend, implying a decoupling of the U-series disequilibria from the major element systematics. The origin of the ‘local’ trend at oceanic ridges is still not resolved (Klein & Langmuir, 1987; Langmuir et al., 1992), but a recent model has suggested that it can be reproduced by mixing of melts derived from episodes of both fractional melting and equilibrium porous flow with the same melt column (Asimow, 1999). In this case, the dynamics of melting and melt extraction is markedly different between the ELSC and CLSC despite the similar spreading rates.

Compared with the Central Lau Basin lavas, samples from the Mariana Trough back-arc basin have significantly higher Na8·0 and the axial ridges are deeper. From the global models for ocean ridges, this would suggest lower overall degrees of melting and a cooler mantle beneath the Mariana Trough (Klein & Langmuir, 1987; Langmuir et al., 1992). Theoretical modelling by Hirschmann et al. (1999) predicts that the melt production per increment of added water is sensitive to initial mantle temperature. Thus, the relationship between F (degree of melting) and source H2O beneath the Central Lau Basin should be different from that determined for the Mariana Trough (Stolper & Newman, 1994), and a higher F for a given H2O content would be expected.

Mantle domains and wedge dynamics beneath the Lau Basin–Tonga arc system

The change in the sense of 238U–230Th fractionation (Fig. 3) in magmas erupted at spreading centres within the Lau Basin coincides with a major shift both in the mantle wedge source composition (Fig. 4) and in the inferred overall degree of melting (Fig. 8). The poorly sampled region between 20° and 21°S on the ELSC is clearly a key transition zone, and this ridge section should be the target of future detailed sampling and geochemical studies within the Lau Basin. The compositional variations along the Valu Fa–ELSC ridge cannot simply be due to a progressive decline to the north in the influence of a fluid component derived directly from the subducting slab, as there is also a significant change in mantle source from ‘Pacific’-type MORB mantle at the Valu Fa Ridge in the south to an ‘Indian’-type MORB mantle beneath the northern part of the ELSC and CLSC. These two different mantle types do not appear to interact with each other beneath the Lau Basin (Hergt & Hawkesworth, 1994), as can be seen from the Pb isotope data (Fig. 4), and thus appear to form isolated flow regimes that will have an important influence on the dynamics of mantle flow and convection within the wedge. A similar situation is seen in other places where these two major mantle domains are found in close proximity, such as at the Australian–Antarctic Discordance (Klein et al., 1988) and in the Vanuatu arc region (Crawford et al., 1995; Peate et al., 1997; Turner et al., 1999).

Seismic tomographic studies in the Central Lau Basin–Central Tonga arc region provide several important observations related to mantle wedge dynamics (Zhao et al., 1997). Beneath the Central Tonga arc, a slow velocity anomaly runs parallel to the dipping slab surface to ∼140 km depth, and this anomaly probably represents the source region for the arc magmas, with melting triggered by input of volatiles from the adjacent slab. Very slow velocity anomalies (5–7%) are found to a depth of ∼100 km beneath the active CLSC and ELSC ridges, which correspond to the region of back-arc magma generation. Below this, moderately slow velocity anomalies (2–4%) extend down to depths of at least 400 km, which suggests that the geodynamics of back-arc spreading might involve deep-seated processes such as deep dehydration reactions in the subducting slab or convective upwelling within the wedge (Zhao et al., 1997). The seismically slow regions beneath the Central Lau Basin spreading centres and beneath the Tonga arc appear to be separated at shallow levels (<100 km). This is consistent with the geochemical observations presented in this paper, which imply that the geodynamic regime in the mantle is distinct between the two regions. These relationships are illustrated on the schematic cross-sections in Fig. 12. The Valu Fa Ridge is propagating very close to the arc front and taps the same ‘Pacific’-type mantle wedge source feeding the Tonga arc volcanoes that has been modified by fluids released by shallow dehydration of the slab. The CLSC and ELSC ridges are associated with a different convection regime within the wedge that involves an ‘Indian’-type mantle source. This source has been modified by the minor addition of compositionally different slab-fluid component derived either from deep dehydration of the slab or from extensive lateral fluid transport from the arc-front region.

SUMMARY

  1. New mass spectrometric U-series disequilibria data on glasses from the back-arc Lau Basin (SW Pacific), the first such study on back-arc magmas, indicate that the sense of 238U–230Th fractionation varies with the distance of the spreading centres behind the trench. Samples from the Valu Fa Ridge close to the arc front all have relatively high (238U/232Th) of 1·28–1·37, and ‘arc-like’ (230Th/238U) <1. In contrast, samples from the Central Lau Basin (ELSC, ILSC and CLSC) to the north, >250 km behind the trench, all have lower (238U/232Th) of 0·98–1·15, and ‘MORB-like’ (230Th/238U) ≥1.

  2. The mantle source of the Valu Fa Ridge magmas is broadly similar to that beneath the Central Tonga arc volcanoes, but with less addition of a compositionally similar slab-derived fluid. 230Th–238U data on Valu Fa samples can therefore be combined with published analyses from the Central Tonga arc to infer a fluid addition event at ∼50 ka, which is consistent with the age of 60 ka estimated from 231Pa–235U analyses of Tonga arc samples (Bourdon et al., 1999). The similar sources and time scales for fluid transfer beneath the Valu Fa Ridge and beneath the arc itself are consistent with a model in which the Valu Fa Ridge is propagating into the arc-front region.

  3. The mantle source beneath the Central Lau Basin is similar in both trace element and isotope composition to Indian MORB mantle, and there is no evidence for a slab-derived component related to the adjacent subduction zone in CLSC magmas. However, a water-rich component is detectable, to a variable degree, in ELSC magmas closer to the arc, and is thus probably related in some way to active subduction processes. The composition of this fluid is different from that found at the Valu Fa Ridge and the Tonga arc-front volcanoes, and this probably reflects a greater equilibration with mantle material during percolation through the wedge.

  4. Melt generation at the Central Lau Basin spreading centres is dominated by normal ‘ridge’-type processes. Samples all have (230Th/238U) ≥1, which indicates initiation of melting in the presence of residual garnet, as is seen in virtually all global MORB. The Central Lau Basin data plot below the broad negative correlation between (230Th/238U) and axial ridge depth shown by global MORB (Bourdon et al., 1996b). The reasons for this are not clear: the effect of increased melt productivity as a result of the water-rich source is probably most likely, although a faster upwelling rate cannot be ruled out.

  5. The shift from 238U excesses to 230Th excesses in Lau back-arc magmas cannot be explained by less addition of a U-rich slab-derived fluid, and it is not just a simple function of increasing distance behind the arc. Different mantle sources are involved beneath the Central Lau Basin (‘Indian’-MORB-type mantle) and the Valu Fa Ridge (‘Pacific’-MORB-type mantle), which are linked to different convective upwellings within the mantle wedge. The transition region between 20° and 21°S on the ELSC is clearly an important target to look closely at the interaction between these different dynamic regimes.

*Corresponding author. Present address: Danish Lithosphere Centre, Øster Voldgade 10-L, DK-1350 Copenhagen K, Denmark. Telephone: +45-38-14-2664. Fax: +45-38-10-0878. E-mail: dwp@dlc.ku.dk

†Present address: GEOMAR Research Centre for Marine Geosciences, Wischhofstrasse 1–3, D-24148 Kiel, Germany.

‡Present address: Department of Geology, University of Bristol, Bristol BS8 1RJ, UK.

We are grateful to Wolfgang Bach and Tracy Vallier for generously providing additional samples for this study, and to Sally Newman (Caltech) for sharing FTIR H2O results on the CD33 glasses with us before publication. We thank Mabs Gilmour and Louise Thomas for their assistance in the lab, and Pete Evans and Rob Hughes for running many of the Ra samples. We are grateful to Simon Turner and Ian Parkinson for numerous stimulating discussions on arc magmatism, and to Wolfgang Bach, S. Fretzdorff and an anonymous reviewer for their constructive reviews of the manuscript. This work was supported by the BRIDGE initiative of NERC (grant GST/02/1161). Isotope research at the Open University is partly funded by NERC. D.W.P. acknowledges support during manuscript preparation from the Danish Lithosphere Centre, which is funded by the Danish National Research Foundation.

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